Observing Systems for Atmospheric Composition: Satellite, Aircraft, Sensor Web and Ground-Based Observational Methods and Strategies

Observing Systems for Atmospheric Composition

Guido Visconti Mark Schoeberl Piero Di Carlo Andreas Wahner William H. Brune Editors

Observing Systems for Atmospheric Composition Satellite, Aircraft, Sensor Web and Ground-Based Observational Methods and Strategies

With 124 Illustrations

Editors Guido Visconti CETEMPS-Dipartimento di Fisica Università degli Studi di L’Aquila via vetoio 67010 Coppito (AQ), Italy [email protected]

Mark Schoeberl Earth Sciences Directorate Mail Code 900 NASA Goddard Space Flight Center Greenbelt, MD 20771, USA [email protected]

Piero Di Carlo CETEMPS-Dipartimento di Fisica Università degli Studi di L’Aquila via vetoio 67010 Coppito (AQ), Italy [email protected]

Andreas Wahner Forschungszentrum Jülich GmbH ICG-II: Troposphäre D-52425 Jülich Germany A. [email protected]

William H. Brune Department of Meteorology Penn State University 505 Walker Building University Park, PA 16802, USA [email protected]

Library of Congress Control Number: 2005938488 ISBN-10: 0-387-30719-2 ISBN-13: 978-0387-30719-0

e-ISBN-10: 0-387-35848-X e-ISBN-13: 978-0387-35848-2

© 2007 Springer Science+Business Media, LLC All rights reserved. This work may not be translated or copied in whole or in part without the written permission of the publisher (Springer Science+Business Media, LLC, 233 Springer Street, New York, NY 10013, USA), except for brief excerpts in connection with reviews or scholarly analysis. Use in connection with any form of information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed is forbidden. The use in this publication of trade names, trademarks, service marks, and similar terms, even if they are not identified as such, is not to be taken as an expression of opinion as to whether or not they are subject to proprietary rights. 9 8 7 6 5 4 3 2 1 springer.com

Preface The goal of the third International Summer School on Atmospheric and Oceanic Science (ISSAOS 2004) was to bring together experts in observing systems and the atmospheric sciences to discuss the need for an observing system for atmospheric composition, its components, and the integration of components into a system. Much of the lecture material was conceptual, with the idea to provide attendees with a context to put their own component of the observing system. The Local Committee, Guido Visconti and Piero Di Carlo, started to think about this school in the summer 2002 and asked William Brune to be a director. Prof. Brune accepted with enthusiasm and was able to get Mark Schoeberl and Andreas Wahner as co-directors. Because the director soon realized that they needed another year to put together all the speakers he had in mind, the school was held from 20–24 September 2004 in L’Aquila (Italy). The speakers were P. K. Bhartia, W. Brune, J. Burrows, J.-P. Cammas, K. Demerjian, H. Fischer, D. Jacob, P. Newman, K. Reichard, V. Rizi, M. Schoeberl, M. Schultz, U. Schumann, A. Thompson, C. Trepte, A. Wahner. This edition of ISSAOS, for the first time, asked students to provide an evaluation of the school at its conclusion. The students generally liked the school, including the quality of the lectures, the opportunities to ask the lecturers questions, and accessibility of the lecturers for conversations. These results encouraged us to put together the lectures of the school in a book to give a larger audience the opportunity to learn about the observational and modeling techniques used to understand the atmospheric composition from satellites, aircraft, and ground-based platforms. For many lectures were two common themes: the role of each component in an observing system for atmospheric composition, and the advances necessary to improve the understanding of atmospheric composition. For the school’s organization we acknowledge the financial contribution of the Italian Ministry of the Environment, and of Center of Excellence CETEMPS. We are also grateful to Simona Marinangeli and Manuela Marinelli for the help in the school’s organization and the hard work of rearranging the lectures in an editorial format. January 2006

G. Visconti P. Di Carlo W. H. Brune M. Schoeberl A. Wahner (Editors)

Contents Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v Introduction I . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 Observing Systems for Atmospheric Composition . . . . . . . . . . . . . . . . . . . . 1 W.H. Brune (Pennsylvania State University, USA) Introduction II . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3 Needs for Sampling on Short Time and Spatial Scales . . . . . . . . . . . . . . . . . 3 W.H. Brune (Pennsylvania State University, USA) Part I: Observations by Satellites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 Chapter 1 Air-Quality Study from Geostationary/High-Altitude Orbits . . . . . . . . . . . . 23 P.K. Bhartia (NASA GSFC, USA) Chapter 2 Aerosol Forcing and the A-Train . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . C. Trepte (NASA LRC, USA)

38

Chapter 3 Total Ozone from Backscattered Ultraviolet Measurements . . . . . . . . . . . . 48 P.K. Bhartia (NASA GSFC, USA) Chapter 4 The EOS Aura Mission . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . M.R. Schoeberl (NASA GSFC, USA)

64

Chapter 5 MIPAS experiment aboard ENVISAT . . . . . . . . . . . . . . . . . . . . . . . . . . . . . H. Fischer (University of Karlsruhe, Germany)

71

Part II: Aircraft and Ground-Based Intensive Campaigns . . . . . . . . . . 83 Chapter 6 Probing the Atmosphere with Research Aircraft-European Aircraft Campaigns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 U. Schumann (DLR Oberpfaffenhofen, Germany)

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Contents

Chapter 7 MOZAIC -Measuring tropospheric constituents from commercial aircraft . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.-P. Cammas (LA/CNRS, France)

97

Chapter 8 Uninhabited Aerial Vehicles: Current and Future Use . . . . . . . . . . . . . . . . . 106 P.A. Newman (NASA GSFC, USA) Chapter 9 U.S. Ground-Based Campaign -PM Supersite Program . . . . . . . . . . . . . . . 119 K.L. Demerjian (University of Albany, USA) Part III: Ground-Based Networks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 Chapter 10 Ozone from Soundings: A Vital Element of Regional and Global Measurement Strategies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 131 A.M. Thompson (NASA GSFC, USA) Chapter 11 LIDAR Networks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 V. Rizi (University of L’Aquila, Italy) Chapter 12 U.S. Federal and State Monitoring Networks . . . . . . . . . . . . . . . . . . . . . . . 159 K.L. Demerjian (University of Albany, USA) Chapter 13 Autonomous Systems and the Sensor Web . . . . . . . . . . . . . . . . . . . . . . . . . 169 K.M. Reichard (Pennslyvania State University, USA) Chapter 14 Comparison of Measurements – Calibration and Validation . . . . . . . . . . . . 182 P.A. Newman (NASA GSFC, USA) Part IV: Output of the Observational Web . . . . . . . . . . . . . . . . . . . . . . . 201 Chapter 15 The Sensor Web: A Future Technique for Increasing Science Return . . . . . 203 M.R. Schoeberl and S. Talabac (NASA GSFC, USA) Chapter 16 Fundamentals of Modeling, Data Assimilation, and High-Performance Computing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 207 R.B. Rood (NASA GSFC, USA)

Contents

ix

Chapter 17 Inverse Modeling Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 230 D. Jacob (Harvard University, USA) Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 239

List of Contributors Dr. Pawan K. Bhartia NASA Goddard Space Flight Center Mail Code 916 Greenbelt, MD 20771 USA e-mail: [email protected] Prof. William H. Brune The Pennsylvania State University Department of Meteorology 505 Walker Building University Park, PA 16802 USA e-mail: [email protected] Dr. Jean-Pierre Cammas CNRS (UMR 5560) Laboratoire d'Aérologie OMP, 14 Avenue E. Belin 31400 - Toulouse France e-mail: [email protected] Prof. Kenneth L. Demerjian State University of New York at Albany Atmospheric Sciences Research Center 251 Fuller Road Albany, NY 12203 USA e-mail: [email protected]

Prof. Herbert Fischer Institut für Meteorologie und Klimaforschung Forschungszentrum Karlsruhe GmbH Hermann-von-Helmholtz-Platz 1 D-76344 Eggenstein-Leopoldshafen Germany e-mail: [email protected] Prof. Daniel J. Jacob Harvard University Pierce Hall, 29 Oxford St. Cambridge, MA 02138 USA e-mail: [email protected] Dr. Paul A Newman NASA-Goddard Space Flight Center Mail Code 916 Building 33, Room E320 Greenbelt, MD 20771 USA e-mail: [email protected] Dr. Karl M. Reichard The Pennsylvania State University Applied Research Laboratory P.O. Box 30 State College, PA 16804 USA e-mail: [email protected]

xii

List of Contributors

Dr. Vincenzo Rizi CETEMPS-Dipartimento di Fisica Università degli Studi di L’Aquila Via vetoio 67010 Coppito-L’Aquila Iytaly e-mail: [email protected]

Prof. Ulrich Schumann Institut fuer Physik der Atmosphaere DLR Oberpfaffenhofen Postfach 1116 D-82230 Wessling Germany e-mail: [email protected]

Dr. Richard B. Rood Earth and Space Data Computing Division Mail Code 930 NASA-Goddard Space Flight Center 8800 Greenbelt Road Greenbelt, MD 20771 USA e-mail: [email protected]

Prof. Anne M. Thompson The Pennsylvania State University Department of Meteorology 510 Walker Building University Park, PA 16802 USA e-mail: [email protected]

Dr. Mark Schoeberl Earth Sciences Directorate Mail Code 900 NASA Goddard Space Flight Center Greenbelt, MD 20771 USA e-mail: [email protected] Dr. Martin Schultz Biogeochemical System Department Max-Planck-Institut für Meteorologie Bundesstr. 55 D-20146 Hamburg Germany e-mail: [email protected]

Dr. Charles “Chip” R. Trepte NASA Langley Research Center Hampton, VA 23681-0001 USA e-mail: [email protected]

Introduction I Observing Systems for Atmospheric Composition WILLIAM H. BRUNE

The concept of observing systems is not new. Even at the dawn of weather satellites in the mid-1960s, planners were conceiving ways to combine the new cloud pictures from satellites with information from radiosondes, aircraft, and surface sites to improve weather forecasting. A 1965 drawing from the United States National Oceanic and Atmospheric Administration illustrates such an observing system.

FIGURE 1. Schematic of a fully integrated environmental monitoring system, circa 1965. (From the United States National Oceanic and Atmospheric Administration image archive.) 1

2

William H. Brune

Even observing systems for atmospheric composition aren’t new. Ozonesondes, aircraft, ground-based networks, field-intensive studies, and lidars have often been used together to advance the understanding of atmospheric composition. Satellites have been part of the observing system for stratospheric composition related to ozone and ozone changes for more than 30 years, but typically not part of the observing systems for atmospheric composition in the troposphere, the lowest layer of Earth’s atmosphere. So, why hold an International Summer School of Oceanic and Atmospheric Sciences on observing systems for atmospheric composition in 2004? We had three main reasons. First, the advances in computer technology, data distribution, collection, and assimilation, and modeling enable the use of information from observing systems in ways that were not possible before. Second, new satellite instruments can make global observations of key atmospheric constituents in the troposphere. Third, with the new computational and observing capabilities, understanding current global atmospheric composition and being able to predict future global atmospheric composition is becoming more of a reality. Our planet is changing. A changing atmospheric composition is coupled with changes in other Earth components to determine the consequences of those changes. An observing system for atmospheric composition will thus help improve the understanding of and predictability for these other Earth components—climate variability and change, the water cycle, human contributions and responses, land use and land cover changes, ecosystems, and the carbon cycle— as well.

FIGURE 1. Schematic of the climate system (from Our Changing Planet, the U.S. Climate Change Science Program for Fiscal Years 2004 and 2005).

Introduction II Needs for Sampling on Short Time and Spatial Scales WILLIAM H. BRUNE

Introduction While the major constituents of the Earth’s atmosphere have not changed much since the advent of humans, the minor components have varied dramatically. Humans and their activities and natural processes both exert a powerful influence on atmospheric composition, sometimes with serious consequences. Understanding the variability and mitigating these consequences require an observing system for atmospheric composition. Components of an observing system for atmospheric composition have existed for decades, although recent advances in sensor technology and computational power make an integrated observing system more realizable. The observing system for atmospheric composition consists of instruments, models, and research. The observing system is needed to develop an understanding of atmospheric composition and the processes that drive it and to provide the capability to confidently predict the interactions between atmospheric composition and changes and Earth and human activities. Two aspects of our existence are particularly affected by variations in atmospheric composition: – air quality (human health, infrastructure and visibility degradation, ecosystem damage) – climate (Earth’s radiative balance and its cascading effects)

Thus, we are interested in the atmospheric composition that influences either atmospheric chemistry or Earth’s radiation balance One way to examine the observing needs for tropospheric composition is the list the issues affecting humans and the atmospheric processes that determine tropospheric composition (Figure 1). The issues range from the global scale to the regional and local scale: climate, atmospheric oxidation, global pollution, carbon balance, regional pollution and haze, and urban pollution and PM2.5 (particles < 2.5 µm in diameter). The processes are not all chemical: emissions 3

4

William H. Brune Processes

Emissions Meteorology Microphysics Chemistry Transport

Issues Climate Sunlight Atmospheric Oxidation

NO

ozone

Global Pollution

HO2

OH

aerosols Carbon Balance

VOCs

Regional Pollution / Haze Urban Pollution / PM2.5

FIGURE 1. Schematic the processes and issues of atmospheric composition.

of atmospheric constituents from surface and airborne sources, meteorology, microphysics, chemistry, and transport. We seek to answer two questions: ● ●

How well can we describe these processes? Can we develop predictive capability for these issues?

The focus of this lecture will be on tropospheric composition and its behavior on short times scales and local to regional spatial scales. Of interest are water vapor (H2O), carbon dioxide (CO2), the pollutants ozone (O3), and small particles and their precursors (particularly nitrogen oxides (NOx=NO+NO2), volatile organic compounds (VOCs), carbon monoxide (CO), sulfur dioxide (SO2), and ammonia (NH3), as well as acids and nitrates, toxic and carcinogenic gases. These atmospheric constituents can be either anthropogenic or natural. Even if we are only interested in pollution, we need to understand the interactions of pollutants with the environment. Consider an atmospheric constituent that is a pollutant (e.g., an atmospheric component that has serious negative consequences for humans and their environment in trace amounts). Predicting the amount of this pollutant, which we will call A, requires observed variables, laboratory measurements of rate coefficients, and computer modeling. Take the example of a simple box model. The time rate of change for an atmospheric constituent, A, can be determined from the equation: qA vdA d[A] = — u ([A] – [A]) —— [A] + —– + P(A) – L(A)[A] – —– o dt ∆x H H

(1)

Introduction II

5

where [A] is the concentration of A (molecules cm−3) qA is the emission rate (molecules cm−2 s−1) H is the height of the mixing box (cm) P(A) is the chemical production rate of A (molecules cm−3s−1) L(A) is the 1st-order loss rate of A (s−1) vdA is the deposition velocity of A (cm s−1) u is the horizontal wind velocity (cm s−1) ∆x is the width of the box in the wind direction (cm) [A]o is the background value for [A] (molecules cm−3) At any given time, any one of these terms can be dominant, depending of the identity of A and the local meteorology. Often, several are comparable. The amount of a pollutant that has adverse effects on humans and their environment depends on the atmospheric constituent and has changed over time as medical and public health studies improve. At present, the United States (EPA, 2004; WHO, 2000) and the World Health Organization have six pollutants whose adverse levels have been established (Table 1). Generally the ambient air-quality standards of the US, WHO, European Union, and other countries are similar, although there are differences. Of these pollutants, CO, SO2, and lead are primary emissions, NO2 (nitrogen dioxide) is rapidly exchanged with NO (nitric oxide), which is a primary emission, particulate matter is both a primary emission and created by gas-to-particle conversion, while ozone is purely a secondary emission generated by atmospheric chemistry. The emissions of greatest interest for local-to-regional scales and times of less than a month or so are those species with lifetimes less than that of CO (Figure 2). A goal of an observing system for atmospheric composition is to develop the capability to link species like CO, O3, NOx (=NO+NO2) to the atmospheric radicals OH, HO2, CH3O2, and NO3, which are the driving forces behind the atmospheric chemistry that creates secondary pollution and modifies primary pollution.

TABLE 1.1. Asian Pollutant Emissions (Gg in year 2000) & 2σ Uncertainties Country China Pollutant SO2 20.400 NOx 11.350 CO2* 3.820 CO 115700 CH4 38.360 NMVOC 17.430 BC 1.050 OC 3.390 NH3 13.570 *

Japan % 13 23 16 156 71 59 484 495 53

800 2.200 1.200 6.810 1.140 1.920 53 74 350

CO2 in Tg. From Streets et al., 2003

India % 9 19 7 34 52 35 83 181 29

5.540 4.590 1.890 63.340 32.850 10.840 600 2.840 7.400

Asia total % 26 48 33 238 67 149 359 544 101

34.300 26.770 9.870 278560 106820 52.150 2.540 10.420 27.520

% 16 37 31 185 65 130 364 450 72

6

William H. Brune CFCs N2O CH4

100 yr CH3CCl3 10 yr CH3Br

Time Scale

1 yr 1 mo

PM2.5 trop O3 SO2

HOOH

CO

NOx 1 day

DMS propene

isoprene

1 hr 100 s

CH3O2

HO2 NO3 1s

1m

OH

100 m

1 km

100 km

1000 km

10,000 km

Spatial Scale

FIGURE 2. Spatial and temporal scales of atmospheric constituents.

Intensive studies show that the understanding of the processes that control atmospheric composition is reasonably good for some conditions. However, it does not yet enable the confident predictability needed for decision making regarding improved air quality, other socio-economic factors, and air-quality improvement deadlines. We examine the current understanding of each process in Figure 1.

Emissions Governmental environmental agencies in the US and Europe have gone to great lengths to quantify a wide range of emissions. The result is that emission inventories are probably best in these two regions. In other regions the emission inventories are less well known. For instance, the inventory of gaseous and primary aerosol emissions in Asia in the year 2000 is fairly certain for atmospheric constituents like SO2, NOx, and CO2, but is far less certain for CO,

Introduction II

7

TABLE 1.2. Emissions from Two Chemical Plants Near Houston, Texas Plant Sweeney

Freeport

Emission inventory observed obs/inventory inventory observed obs/inventory

NOx (kmole/hr) 14 15 1.07 31 30 0.97

C2H4 / NO (mole/mole) 0.01 3.6 360 0.03 1.5 50

C3H6 / NO (mole/mole) 0.01 2.0 200 0.01 0.5 50

nonmethane volatile organic compounds (NMVOCs), and organic carbon (OC) and black carbon (BC) aerosols (Table 2). Even in U.S. urban areas, inventories of some emissions are found to be significantly in error. An example is Houston, Texas, in September 2000, during the TexAQS2000 intensive field campaigns (TexAQS2000). The NOAA P-3 sampled plumes from two petrochemical plants on the coast south of Houston. In the plumes were highly elevated levels of ethane (C2H4) and propene (C3H6). By applying a transport model to the plume measurements, the emission rates were derived. The resulting emissions of NOx agree with the emissions inventory (Table 3). However, the resulting ratios of ethene and propene to NOx are more than 50–200 times larger than expected from the emissions inventory. The reasons for this difference are not clear, but it could be leaks or many small spillages that are less than the level that is required to be reported. What is clear is that the emissions inventories drastically underestimated the actual alkene emissions. This error in the emissions inventory for ethene and propene is important for Houston’s pollution chemistry. The hydroperoxyl radical, HO2, is a precursor atmospheric constituent that leads to ozone production. It is mostly made during the day by processes involving sunlight, but is sometimes made at night, when ozone reacts with alkenes like propene and ethene. In Houston, a large nighttime propene spike leads to an HO2 spike that is as large as HO2 gets during the day (Figure 3). Thus, emission errors will propagate into errors in the modeling of radical chemistry and ozone production. Conclusions regarding emissions: ●

●

●

●

Some U.S. emissions are known quite well (e.g., NOx from U.S. power plants); others are known quite poorly (e.g., ethene and propene). In the developing world, emission inventories are generally even less well known. There is also a factor of 2–3 uncertainty in regional natural emissions (such as for isoprene, a highly reactive VOC emitted by trees). Since emissions are the fuel for tropospheric (and stratospheric) photochemistry, it is important to know well the emissions, their distributions, and variations.

8

William H. Brune 60 HO2 (pptv) Propene (ppbv)

50

Mixing ratio

40

30

20

10

0 midnight

6 am

noon

6 pm

midnight

FIGURE 3. Propone emissions in the presence of ozone initiate fast atmospheric chemistry, as evidenced by the nighttime spike in propene and the reactive free radical HO2 during the TexAQS2000 intensive study.

Tropospheric Chemistry Where does ozone come from? In the stratosphere, ozone can result from the direct dissociation of molecular oxygen by ultraviolet light into oxygen atoms, which then join with molecular oxygen to form ozone. Ozone formation in the troposphere is more complex. It requires the presence of ultraviolet sunlight, although not as energetic as in the stratosphere, nitrogen oxides (NOx = NO + NO2), and volatile organic compounds (VOCs), which are the source of the free radicals, hydroxyl (OH), and hydroperoxyl (HO2). Ozone is a secondary pollutant—its only source is in the atmosphere. It is lost, though, both on surfaces and by photochemical reactions, as given in the following equation: d[O3] vdO3 u ([O ] – [O ]) —— = P(O3) – L(O3)[O3] – —– [O3] + —– 3 o 3 dt ∆x H where P(O3 ) = kHO2 + NO [NO][HO2 ] + ∑ kRO2i + NO [NO][RO2i] i

L(O3 ) = J(O3)fH2O + kHO2 + O3 [O3][HO2] + kOH + O3 [O3][OH]

Introduction II

9

Ozone losses include its photodestruction, followed by reaction with H2O to form OH, direct reaction with OH and HO2, and deposition on surfaces. Deposition and formation of OH are often the two largest ozone destruction processes. The instantaneous ozone production, P(O3), comes from a reaction that cycles HO2 to OH: HO2 + NO → OH + NO2 NO2 + sunlight → NO + O O+ O2 + M → O3 + M Ozone is also made by replacing HO2 in the chemical equation above with RO2, where R is CH3, C2H5—a hydrocarbon with an odd number of hydrogen atoms. The instantaneous ozone production rate is determined to some extent by the cycling of HOx between OH and HO2 where HOx = OH + HO2 (Figure 4). Most HOx is produced by photochemical processes as OH, although in environments where formaldehyde (HCHO) is high, a significant fraction of HOx can come from the production of HO2 by HCHO destruction by sunlight. Once created, OH is rapidly cycled to HO2 by OH reactions with CO, CH4, VOCs, and O3 while HO2 is cycled to OH by HO2 reactions with NO and O3. If the sources of OH and HO2 were cut off, then OH would cycle to HO2 in about a second. The reaction of OH with VOCs typically produces RO2 before it produces HO2. As a result, RO2 abundances are roughly equal to HO2 abundances. The amount of RO2 and HO2 produced is thus related to P(HOx) and the amount and type of VOCs that react with OH. When NO is greater than a few 10’s of pptv, the reaction of HO2 with NO dominates the reaction of HO2 with O3 and ozone is produced, not destroyed. Eventually, HOx reacts to form more stable atmospheric constituents and the cycle is terminated. The reaction products depend on the HOx P(HO2)

P(OH)

CO, CH4, O3,VOCs HO2

OH

RO2 NO, O3

increasing NOx

Peroxides (HOOH) (CH3OOH)

H2O HO2NO2 O3

acids (HNO3)

FIGURE 4. A schematic of the photochemistry of OH and HO2 as a function of NOx.

10

William H. Brune

production rates (P(OH) and P(HO2)), and on the NOx abundances (Figure 4). Consider a case where P(HOx) and the CO, CH4, and VOC emissions are constant. At low NOx, the HOx is shifted to HO2 (the HO2/OH ratio is large) and the reactions of HO2 with HO2 and RO2 produce the peroxides HOOH and ROOH. When NOx is greater, more HOx is shifted to OH (the HO2/OH ratio is less) and the reaction of HO2 with OH to form H2O slightly dominates. When NOx is very abundant, such as in a more polluted region or in the upper troposphere and stratosphere, then HOx is shifted even more toward OH and the reaction of OH with NO2 to form HNO3 is most important. A similar diagram can be drawn for NOx. In this case, the primary cycling between NO and NO2 occurs by the reaction sequence: NO2 + sunlight → NO + O O + O2 + M → O3 + M NO + O3 → NO2 + O2 NO + HO2 → NO2 + OH NO + RO2 → NO2 + RO The first two equations show how NO2 is converted into ozone. However, the reaction of NO with O3 destroys O3; a steady-state cycle forms during the day. No new O3 is really created by this cycle. Instead, the O atom is merely exchanged between O3 and NO2. However, the reaction of NO with HO2 and RO2 does make new O3. The coupling of the HOx and NOx cycles produces O3 (Figure 5). As the NOx abundance is increased, OH first rises as more HOx is shifted from HO2 to OH and then falls, as OH is rapidly converted to HNO3. HO2 is relatively unaffected by increasing NOx until NOx reaches about 1 ppbv; then HO2 decreases increasingly rapidly as HOx is shifted from HO2 to OH and as OH (and thus HOx) is removed by NO2. The resultant instantaneous ozone production rate thus increases until NOx reaches a few ppbv and then decreases. Thus, instantaneous ozone production exhibits significant nonlinear behavior as a function of NOx.

HO2 10

1

0.1

0.1

P(O3)

OH

1

10

FIGURE 5. Variation of OH, HO2, and P(O3) as a function of NOx.

Introduction II

11

An example of this nonlinear behavior can be seen in observations made during the TexAQS 2000 intensive field campaign in Houston, Texas (Figure 6). The behavior of NO, NO2, and O3 are shown for three days in August. Morning rush hour is marked by peaks in NO and NO2, since NO is emitted by automobile and truck engines. Evening rush hour is also seen, although the peak values are smaller. However, it is clear that O3 is building up over these three days, which means that O3 production is occurring. NOx reaches several 10’s of ppbv during rush hour. Thus, NO and HO2 should anticorrelate. They do (Figure 7). However, since P(O3) ~ [NO] [HO2], it can be large even when a small amount of HO2 reacts with a large amount of NO, and vice versa. P(O3) remains between 40 ppbv/hr and 60 ppbv/hr from the hours of about 8:00 to 14:00. In this period, roughly 300 ppbv of ozone is created, which is proportional to the total ozone observed. Obtaining a more accurate ozone balance requires knowledge not only of the photochemistry, but also of the height of the mixed layer, the ozone loss on Earth’s surface, and the ozone transport to and away from the region of interest. The NO abundance is related to the NOx abundance, and the RO2 and HO2 abundances are related to the VOC abundance. Because NOx and VOCs are the primary emissions that contribute to ozone production, an ozone isopleth plot as a function of VOCs and NOx abundance should provide guidance for the optimum regulatory actions to reduce ozone pollution. An early diagram (Figure 8) met the following conditions: ● ●

It was chosen for a particular region, or box, which could be moving. The time duration of the run was chosen to capture the maximum ozone production.

NO O3 NO2

200

150

100

50

0

0

6

12

18

0

6

12

18

0

6

12

18

0

FIGURE 6. NO, NO2, and O3 during a pollution episode in Houston, Texas, during the TexAQS 2000 air-quality study.

12

William H. Brune 100 NO (ppb) HO2 (pptv) O3 (ppbv) P(O3) (ppb/hr)

80

60

40

20

0

0

6

12

18

FIGURE 7. NO, HO2, and P(O3) during a pollution episode in Houston, Texas, during the TexAQS 2000 intensive field campaign.

●

The entire plot was generated by changing the VOC/NOx mixture while leaving the meteorological conditions constant.

Regions above the ridgeline at VOC/NOx = 8 are “VOC-limited”; more VOCs are need to produce more ozone. Regions on the plot below the ridgeline are “NOxlimited”; more NOx is needed to produce more ozone. Urban areas are typically 0.28 O3 (ppm) = 0.08 0.16 0.24

4 1 0.34

0.40

0.24

NOx (ppm)

VOC-limited (rush hour0.20 urban) 0.12 0.20

0.30 0.28 0.32

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=

8 1

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0.12 0.08

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VOC (ppmC)

FIGURE 8. An example of an ozone isopleth diagram (adapted from NRC, 1990).

2.0

Introduction II

13

“VOC-limited,” while rural and suburban areas and midday urban areas are often “NOx-limited.” This diagram indicates that reducing VOCs will result in reduced produced ozone, while reducing NOx for areas above the ridgeline can result in an increase in ozone until the VOC/NOx ratio drops below 8. This behavior results from the nonlinear coupling between HO2 and NO. Conclusions regarding atmospheric chemistry. ●

●

●

The nonlinear behavior of HOx–NOx chemistry complicates the analysis of ozone production rates from observing systems that have low spatial and temporal resolution. Other products of tropospheric oxidation have effects on human health and climate, such as formaldehyde (HCHO), peroxyacetyl nitrate (PAN). The NO3 radical plays a nighttime role similar to OH’s daytime role.

Particle Microphysics and Chemistry Particles have severe impacts on both human health and climate. Particles can be emitted directly from sources, such as diesel engines, or created in the atmosphere by gas-to-particle conversion. Particles less than 2.5 µm in diameter are particularly important for human health. They are small enough to navigate the respiratory system’s passageways and are deposited deep within the lungs, where they can enter the bloodstream. PM2.5 has been linked to severe cardio-respiratory problems such as chronic asthma and bronchitis and even sudden heart attacks. At the same time, they are also important for Earth’s radiation budget because they are about the size of the wavelength of the incoming solar radiation and are thus efficient at scattering or absorbing light. In gas-to-particle conversion, volatile gases react with OH, O3, or NO3 to form gases that have low vapor pressures. These either condense on pre-existing aerosols or nucleate to form new ones. The main reactions involve SO2, NO2, VOCs, and NH3: SO2 + OH →→ H2SO4 →→ particles NO2 + OH + N2 → N2 + HNO3 →→ particles VOCi + OH → VOCk (low vapor pressure) →→ particles NH3 →→ particles The emission of primary particles and the gas-to-particle conversion of secondary particles lead to particle size distributions, which because of microphysics tend to have certain characteristics (Figure 9). Primary particles tend to be larger than 0.01 µm, while secondary particles are created at sizes in the 0.001–0.01 range when nucleation of low-vapor-pressure gases. Both types of particles grow by coagulation, condensation of both low-vapor-pressure gases and water vapor. These processes create particle size distributions: the nucleation mode (diameter < 0.01 µm), the Aiken mode (0.01 µm < diameter < 0.08 µm), the accumulation mode (0.08 µm < diameter < 2.5 µm), and the coarse mode (d > 2.5 µm). The nucleation and Aiken modes are rapidly lost by coagulation; the coarse mode is

14

William H. Brune Emissions of Primary Particles

Emissions of Gaseous Precursors

In-cloud Reaction Heterogeneous Reactions Coalesc-

Condensation

ence Evaporation In-cloud Reaction

Coagulation

Homogeneous Nucleation Reactions (OH, O3, N03)

Activation and Scavenging

Coagulation Dry Deposition 0.01

0.001

Activation and Hydrosopic Growth

0.1

1

Wet Deposition 10

Particle Diameter Gas Molecules

P

A

R

T

I

C

L

E

S

Cloud drops

FIGURE 9. A schematic of atmospheric particles processes and size distributions (from US CCSP, 2004).

rapidly lost by deposition; the accumulation mode has a lifetime of a few weeks or so. Equally important to the particle size distributions is the chemical composition. While no direct link has yet been made between the chemical composition of PM2.5 to human health, chemical composition has a major impact on climate. For example, some particles, such as sulfate (H2SO4), reflect incoming solar light, while other particles, such as soot from diesel exhaust, absorb incoming solar light. The impact on Earth’s radiation balance is not trivial. Small particles also have indirect effects on Earth’s radiative balance. The nucleation of cloud drops and ice from pure water vapor is unrealistic in Earth’s atmosphere. Instead, nucleation starts with particles. However, only a small subset of atmospheric particles are good cloud condensation nuclei and even a smaller subset are good ice nuclei. The chemical composition, along with the particle size, determines a particle’s ability to be a cloud condensation nuclei (CCN) or an ice nuclei (IN). In the troposphere, particles often contain a combination of sulfate, nitrate, organic or black carbon, and ammonia. The sulfate content in the U.S. East Coast is much greater than it is on the US West Coast, primarily because of the use of sulfur-containing coal in Midwest power plants. Organic content appears to be fairly ubiquitous, even in very remote regions of the Pacific Ocean. Measurements were made with an aerosol mass spectrometer (AMS) in Scotland and Korea (Figure 10). In the Scotland sample, the size distribution is mostly in the Aiken mode and is dominated by the organic and nitrate components. In the Korean sample, the particle size distribution shows a growing accumulation mode and is dominated by the sulfate component. Peaks in the organic

Introduction II 3

15

6 Edinburgh, Scotland November, 2000

4

µg/m3

2

Cheju-Do Island, Korea 5 April, 2001

Organic Nitrate Sulfate

Sulfate Organic Nitrate

3 1

2 1

0

0 456789 2 3 4 5 6 78 9 100 1000 Aerodynamic Diameter (nm)

2

3

5 6 7 89 2 3 4 5 6 7 89 100 1000 Aerodynamic Diameter (nm)

2

3

FIGURE 10. Particle size distributions and chemical composition in a city (Edinburgh, Scotland) and in a remote region downwind of urban and industrial regions (Cheju-Do Island, Korea). (from Hugh Coe, private communication, 2003).

component of the mass spectrum indicate that in Scotland the organic component consists primarily of hydrogen and carbon, while in Korea, the organic component has been oxygenated in the transit from the source regions. This oxidation process alters the particles’ capability to act as CCN. It was generally thought that there were always enough particles present in urban areas that low-vapor-pressure gases would always condense on pre-existing particles and not form new particles. However, recent measurements in Pittsburgh, PA, and other cities have shown that under certain meteorological conditions, such as recent frontal passages, new particles can nucleate (Figure 11). At 9 a.m., OH was created and reacted with SO2 and VOCs to form new particles. This burst of new particles then rapidly coagulated, creating fewer, larger particles in a few hours. That new, small particles are formed may have implications for human health. Conclusions for microphysics. ● ●

●

● ●

Airborne particles participate in both air quality and Earth’s radiative balance. Particle distribution usually consists of particles emitted directly from sources and particles created by gas-to-particle conversion. Particles’ roles in human and health and climate are determined by both the particle size distribution and the particles’ chemical composition Particles containing organics are widespread. PM2.5 is a major research thrust—for reasons of both health effects and climate.

Meteorology Meteorology affects the distribution of atmospheric constituents and atmospheric composition on spatial scales from turbulence to global transport and on temporal scales from seconds to decades. Without a knowledge of meteorology, it is

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William H. Brune

Particle diameter (nm)

Shenley Park 11 August 2001 5.0 500

4.0

100

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10 2.0 µg/m3

120,000 80,000 40,000 0 0

3

6

9

12 15 Time (EDT)

18

21

24

FIGURE 11. New particle formation & growth, 11/08/2001, Pittsburgh, PA. (From S. Pandis, private communication, 2003).

impossible to understand measurements of almost all atmospheric constituents, with the possible exception of the short-lived radicals CH3O2, NO3, HO2 in most conditions and OH in essentially all conditions (Figure 2). The OH lifetime is a second or less; its abundance and behavior are defined exclusively by reactions involving NOx, VOCs, CO, and a few other constituents, and P(HOx). However, for OH, solar photolysis must be well known, and that depends on clouds. The interaction between meteorology and atmospheric composition continues to be an active research topic. To illustrate the impact of meteorology on atmospheric composition, and its measurement, consider the planetary boundary layer, or PBL (Figure 12). During the day, surface heating generates convection, which raises and mixes surface gaseous and particulate emissions throughout the convective mixed layer, which is typically 1 to 4 km high and capped by a cloud layer. The circulation time constant is about an hour. While PBL constituents are vented into the free troposphere, especially by convective systems, the PBL can often be considered to be a box into which the surface emissions are mixed. As sunset approaches, convection becomes less active and a lower, stable PBL forms. Atmospheric constituents that were higher in the convective mixed layer when convection ceased remain in the residual layer, disconnected to the surface. Any surface emissions occurring at night remain in the nocturnal boundary layer. During rush hour near sunrise, high levels of vehicle exhaust, including NOx, VOCs, and particulates, build up in the shallow boundary layer. As the sun heats

Introduction II

17

2000 Free Atmosphere Height (nm)

Cloud Layer

Entrainment Zone Capping Inversion Entrainment Zone

1000

Residual Layer Convective Mixed Layer (pollutants can remain, mixing down at sunrise) rush hour Stable (Nocturnal) Boundary Layer rapid cycling

0 Noon

Sunset

Surface Layer Midnight Sunrise

Noon

Local Time

FIGURE 12. The typical behavior of the planetary boundary layer over the course of a day. (Adapted from Stull, Boundary Layer Meteorology, Kluwer, 1988).

the surface and convection begins, the air in the shallow boundary layer is mixed with air from the residual layer. This mixing increases until near midday, when the PBL height is now again greatest. Near the end of rush hour, surface emissions are mixed into a much larger volume, resulting in smaller pollutant mixing ratios since the mixing is into a greater volume. The simultaneous pulse of pollutant emissions at morning rush hour and the increase in the PBL height after sunrise have important consequences. First, because the chemistry that produces ozone and particles is nonlinear with respect to the emitted pollutants’ concentrations, the change in pollutant concentrations due to mixing into an increasing volume affects the production rates of ozone and particles. Second, the behavior of surface-based measurements of pollutant abundances can only be interpreted if the meteorology is known, particularly the variation of the PBL height. The horizontal winds also strongly influence the distribution of pollution and atmospheric constituents, even on a neighborhood scale in urban areas. Chemical transport models transport atmospheric constituents from their emissions sources while transforming them with chemistry and microphysics. Urban airshed models use forecasts and analyses of meteorology on the few kilometer to synoptic scales (~1000 km). Cloud-scale meteorology has two main effects. First, convective lifting of pollutants lifts local and regional pollutants to new heights, where they can be spread and where they can influence the photochemistry. Second, the process of going through clouds can scrub water soluble gases, changing the atmospheric composition and thus the chemistry. Synoptic-scale meteorology creates the conditions necessary for pollution episodes. Stagnation episodes lead to smog (e.g., U.S. East Coast). Stable layers both create and inhibit regional surface pollution (e.g., Los Angeles, CA).

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Synoptic-scale frontal systems lift pollutants and their precursors off the surface and into the free troposphere (e.g., Southeast Asia). Global-scale winds move pollutants great distances from one region to another. There is growing recognition that in the Northern Hemisphere, pollution outflow from one continent is reaching the next continent downwind to the east. Thus, pollution from the United States gets to Europe; pollution from Europe gets to Asia; and pollution in Asia gets to the United States. Countries like Japan that are immediately downwind of industrialized and urbanized regions in China encounter more severe pollution transport. As developing countries in Asia industrialize, global pollution levels are likely to rise, with consequences for regions downwind. Conclusions for meteorology. ● ●

●

Meteorology on all scales influences the atmospheric composition. Because ozone and particle production can depend on constituent concentrations in a nonlinear way, the actual distribution of atmospheric composition can have a large impact on secondary pollutant production. The atmospheric composition of one region is influenced by the atmospheric composition of the upwind regions, even when those regions are thousands of kilometers away. Air quality has become a global issue.

Where to Go from Here? Intensive measurements and monitoring of subsets of processes teach us a lot, but ... ●

●

●

●

●

●

Which emissions, chemistry, microphysics, and meteorology are in common? – We need widespread, long-term process studies. Are observed differences due to differences in specific models and instruments? – We need frequent intercomparisons. Can calculations/observations of linear behavior be extrapolated, especially as the scientific basis for regulatory actions? – We need to understand atmospheric behavior over the entire range of variation; models must match both low and high values. Is “matching” ozone, or OH, or PM2.5 good enough? – We need to match all species (e.g., PAN, NOx, etc.) and meteorology (e.g., PBL height, trajectories) that have importance influence on ozone and aerosols. Are convection and PBL mixing important? – We need to determine influence over the range of transport scales. We can’t do chemistry without doing meteorology! What do we need to do to understand the problem? – We need to learn how to combine measurements from satellites, ground networks, intensive field campaigns, and models on several scales to build an integrated system for creating understanding and predictive capability.

Introduction II

19

References U.S. Environmental Protection Agency (EPA), http://www.epa.gov/air/criteria.html, 2004. National Research Council, Rethinking the Ozone Problem in Urban and Regional Air Pollution, National Academy Press, Washington, DC, 1991. Our Changing Planet, The U.S. CCSP for fiscal years 2004 and 2005, pdf file at http://www.climatescience.gov, 2004. Strategic Plan for the U.S. Climate Change Science Program, pdf file at http://www. climatescience.gov, 2003. Streets, D.G., T.C. Bond, G.R. Carmichael, S.D. Fernandes, Q. Fu, D. He, Z. Klimont, S.M. Nelson, N.Y. Tsai, M.Q. Wang, J-H. Woo, and K.F. Yarber, An inventory of gaseous and primary aerosol emissions in Asia in the year 2000, J. Geophys. Res., 108, D21, 8809, doi:10.1029/2002JD003093, 2003. WHO (World Health Organization), Guidelines for Air Quality. World Health Organization, Geneva, 2000.

Part I Observations by Satellites

Chapter 1 Air-Quality Study from Geostationary/High-Altitude Orbits PAWAN K. BHARTIA

Introduction What determines air quality? There are two ingredients in the air that we breathe that in high concentrations cause poor air quality (AQ). They are particulates (aerosols) smaller than 2.5 µm, called PM2.5, and ozone. In addition, the US Environmental Protection Agency (EPA) regulates sulfur dioxide (SO2), nitrogen dioxide (NO2), and carbon monoxide (CO), which lead to enhanced concentrations of aerosols and O3. SO2 combines with water vapor to form sulfuric acid droplets that are the major constituents of urban smog. In the presence of sunlight, NO2 photolyzes into NO and O; the latter combines with O2 to form O3. CO (as well as some volatile organic compounds) helps recycle NO back into NO2; so the O3 production continues until there is sunlight, or NO is converted by a competing reaction into HNO3 and removed from the system. There are serious health and environmental effects associated with high concentrations of aerosols and O3. Small-size particulates can enter the lung causing breathing problems; they can also deposit cancerous materials, often found on the surface of these aerosols, into the lung. Enhanced concentrations of O3 can affect the immune system and decrease agricultural productivity. However, it is not possible to control the amount of these pollutants simply by controlling the local emission of the source gases, for particulates and long-lived trace gases, such as CO and O3, can be transported to large distances. For example, dust from the Sahara regularly blows north over Europe and is transported west to the Caribbean Islands and to the southeastern US by prevailing winds. Dust from China frequently crosses the Pacific to reach the northwestern US. Pollutants produced by biomass burning in Africa have been detected over some pristine Pacific islands. There are also natural events, such as forest fires, volcanic emission, etc, that can degrade AQ over large areas. Finally, while the AQ is obviously affected by weather and change in climate, there is also a reverse connection. Aerosols have several direct and indirect effects on cloud and radiation. They scatter sunlight and can increase cloud albedo. The carbonaceous aerosols (aerosols containing soot) reduce the amount of sunlight reaching the surface, thus suppressing

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Pawan K. Bhartia

the hydrological cycle. They also warm the atmospheric layer in which they are located and can suppress cloud formation. Thus, the key scientific questions related to air quality one seeks to answer using satellite, ground-based, and aircraft instruments are ● ●

What are the major sources of air pollution? How natural and anthropogenic processes interact in determining global air quality? What are the relative influences of natural and manmade sources? How are the pollutants transported across regions and continents? How do radiation, weather, and climate affect air quality? How does air quality affect weather and climate? ■ ■ ■

●

Measurement of Air Quality by Satellite Instruments The five pollutants mentioned above (O3, SO2, NO2, CO, and PM2.5) are referred to as “criteria pollutants” by the US EPA. It turns out that all five can be measured by remote sensing instruments operating from satellites. In addition, satellite instruments can measure HCHO and HNO3, both of which are related to the reactions that determine air quality, and finally, it is possible to estimate actinic flux in UV and visible, needed for the calculation of O3 and NO2 photolysis rates, from satellites. The satellite instruments can be classified either based on the wavelength range in which they operate or based on their viewing geometry. The limb-viewing instruments, operating from visible to microwave wavelengths, are best for measuring particulates and trace gases in the upper troposphere (and above) with high sensitivity and good vertical resolution. However, in this lecture we will not consider such techniques; neither will we consider nadir-viewing thermal IR techniques for measuring trace gases in the middle and upper troposphere. This is because so far there has been very little experience in deriving AQ-related information from these techniques, except for CO. (This situation is likely to change in the coming months, as the data from the TES instrument, launched on July 15, 2004, on the EOS Aura satellite, are released.). In this lecture we will instead focus on the nadir-viewing techniques employed by the TOMS, GOME, and SCIAMACHY satellite instruments to measure the pollutants in the lower troposphere using reflected/scattered sunlight in UV, visible, and near-IR wavelengths (0.3–2.5 µm). The TOMS measurement series started in November 1978, and the GOME series in July 1995. SCIAMACHY employs techniques similar to GOME but provides higher spatial resolution and measures in the near-IR to retrieve CO column amounts. The following provides overviews of the techniques used by these instruments to measure the 5 criteria pollutants from space.

Aerosol Measurement To study AQ one needs the surface concentration and size distribution of aerosols. In addition one needs their vertical distribution and absorptive properties to do

1. Air-Quality Study from Geostationary/High-Altitude Orbits

25

photolysis calculations. Although the satellites cannot directly measure these quantities, they can provide important information to test, improve, and constrain the chemical and transport models that can provide the needed information. In the future, it may be possible to use data assimilation techniques, in which one statistically combines information in models and measurements, to produce a consistent three-dimensional aerosol field. For decades aerosol optical depth has been derived from the measurement of top-of-the-atmosphere reflectance (TOAR) at a single wavelength in the visible. Figure 1 provides a simplified explanation of how the technique works. (The symbol ρ, commonly used for TOAR, will be defined in Figure 5 of page 53 on the derivation of total O3 from the buv technique. It is a convenient way to convert the radiance measured by a satellite instrument into a dimensionless quantity that behaves like reflectivity.) From Figure 1, in absence of aerosols ρ becomes equal to the surface reflectance R; aerosols scattering (2nd term) increases ρ, while the 3rd and the 4th terms, which represent loss of sunlight to space by aerosol backscattering, and conversion into heat by aerosol absorption, respectively, can make ρ smaller than R for large R. If R is small, one can derive aerosol scattering optical depth (τs) from the 2nd term by assuming the aerosol scattering phase function P(Θ). Figure 1 shows several techniques for assessing P(Θ). The MODIS instrument flying on EOS Terra and Aqua satellites makes multispectral measurements of ρ to determine if the aerosols consist of small or large particles, which helps in the determination of P(Θ). The MODIS also assumes a value for the aerosol single scattering albedo (ratio of τs and τext) to convert τs into aerosol

FIGURE 1. A simplified formula for the top-of-the-atmosphere reflectance measured by satellite instruments at visible/near-IR wavelengths.

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Pawan K. Bhartia

extinction optical depth, τext. Figure 1 shows an estimate of the accuracy in deriving τext from MODIS [Remer et al., 2004]. One notes that the MODIS retrieval accuracy is significantly worse over land; over bright surfaces, e.g., deserts, MODIS is unable to retrieve any aerosol information. (This is because the surface reflectivity of land is more variable, and the last two terms in the equation in Figure 1 reduce the aerosol signal over bright surfaces.). Recently, a new algorithm, called the “deep-blue” algorithm, has been developed [Hsu et al., 2004] to improve the estimation of τext over land. This algorithm uses the 412and 490-nm MODIS channels where the land is darker than at the longer wavelengths. Figure 2 shows the τext and Angstrom coefficient (a rough measure of aerosol size) derived from this algorithm using data from the SeaWIFS instrument, which has similar channels. Finally, we describe a technique to estimate aerosol absorption optical depth τa using measurements of ρ at two ultraviolet wavelengths. The physical basis of this method is shown in Figure 3. Note the formula in Figure 3 is very similar to that in Figure 1, except the surface reflectivity is now replaced with ρ0, the TOAR in absence of aerosols. In ultraviolet, ρ0 is usually much larger than the surface reflectivity and has strong wavelength (λ) dependence caused by Rayleigh scattering from the molecular atmosphere. This leads to a curious

FIGURE 2. Aerosol properties over the Ganges Valley in India derived from the “deepblue” algorithm using SeaWIFS data. Satellite results are compared with a ground-based AERONET station at Kanpur, India. (Figure courtesy of Christina Hsu, UMBC, Maryland, USA.)

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27

FIGURE 3. A simplified formula for the top-of-the-atmosphere reflectance measured by satellite instruments at UV wavelengths.

and a highly nonintuitive result: when τa is nonzero and significant (i.e., when significant amount of UV-absorbing aerosols, such as smoke and desert dust containing iron oxide are present), λ-dependence of ρ0 causes the 4th term to have λ-dependence, which is over and above the λ-dependence of τa itself and modifies the λ-dependence of the 2nd term. For example, for smoke, which consists of small particles, τs increases with decrease in λ, but since τa is nearly independent of λ, the 4th term also does the same with λ, thus ρ−ρ0 can either decrease or increase with λ depending on the relative magnitudes of the various terms. Generally speaking, k3 is nearly zero for aerosols near the surface but increases as aerosols get elevated, so spectral dependence of ρ−ρ0 can change sign depending the height of the smoke. By contrast, τ2 for desert dust is independent of λ, but since τ2 increases with decrease in λ, ρ− ρ0 decreases with λ, irrespective of the height of the dust layer. These features have led to a unique method of detecting UV-absorbing aerosols from spectral contrast of measurements of ρ− ρ~0 at two UV wavelengths. This so-called aerosol index method has been applied to the TOMS data to study the transport of dust and smoke over the globe [Hsu et al., 1996; Herman et al., 1997]. A notable feature of the technique is that, since the aerosol index over cloud is zero, clouds do not show up in the aerosol index maps at all (Figure 4) unless the aerosols are above or embedded in the cloud. Indeed, the aerosol index method is the only (passive) remote sending technique that can detect (UV-absorbing) aerosols over clouds (Figure 5) as well as over very bright snow/ice covered surfaces.

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FIGURE 4. This false color map of TOMS aerosol index (AI) shows smoke from fires in the western US. AI is produced from the spectral contrast of ρ−ρ0 between two TOMS wavelengths. The color of the land in the background is a visible RGB image from another sensor. Note the absence of clouds in the picture; the TOMS AI algorithm is not sensitive to clouds.

FIGURE 5. Right side of the figure shows a true color map of the SE China showing a thick cloud cover. The left side of the figure shows a TOMS AI map over the same region showing the presence of smoke, which is embedded in the cloud or above it. Upon close examination the right side of the figure shows that the clouds under the yellow oval are indeed darker, confirming the TOMS finding.

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Using aerosol models and the Mie scattering theory it is possible to make quantitative estimates of τa and τs from measurements of ρ at two ultraviolet wavelengths [Torres et al., 1998]. With future instruments, the visible, the deep blue, and the UV methods can be combined to make estimates of both the scattering and absorption optical depths of aerosols over both land and ocean. Currently, the largest uncertainty in estimating τa comes from the fact that k3 in Figure 3 is dependent on aerosol vertical distribution. To improve this estimate one needs space-based aerosol lidars.

Estimation of Tropospheric O3 Estimation of tropospheric O3 remains a difficult challenge for satellite instruments, for the radiation that is sensitive to the tropospheric O3 has to pass through the stratosphere before reaching the satellite. The absorption of this radiation due to stratospheric O3 is typically 20 or more times greater than due to tropospheric O3. Moreover, the radiation generated in the troposphere is multiply scattered, which smears out the profile information. Finally, molecular and cloud scattering can significantly reduce the sensitivity to O3 in the lower troposphere. So far there is no evidence that any satellite technique can sense the O3 near the surface. The best that satellite instruments have been able to achieve so far is to estimate the tropospheric O3 column from total column O3 by subtracting the stratospheric O3 overburden. There are several ways of doing so, including the use of limbviewing instruments, such as MLS and SAGE, that can measure down to the tropopause to estimate the stratospheric O3 overburden [Fishman et al., 1990; Ziemke et al., 1998]. Here we describe a novel method, called the “convective cloud difference” (CCD) method in which total O3 measurements made above high altitude clouds are used to estimate the stratospheric overburden [Ziemke et al., 1998]. The principle of the method is illustrated in Figure 6. Based on Kley et al. [1996] the deep convective clouds that often form in the Pacific Ocean pump very clean marine boundary layer air into the upper troposphere making the mixing ratio of O3 in tropical transition layer (TTL), the layer between the top of these clouds and the tropopause, very small. Indeed Kley et al. sometimes detected near-zero ozone mixing ratio in the TTL. This means that the Pacific TTL often contains less than 1 DU of O3, thus providing a very nice way of measuring the stratospheric column O3 by measuring the column O3 above very high deep-convective clouds. Figure 7 shows comparison of the stratospheric O3 column derived using the CCD method with those derived by integrating the O3 profiles measured by the SAGE instrument. SAGE is a solar occultation instrument that provides O3 profiles with 1 km vertical resolution from tropopause to 60 km. Note that due to its limited sampling, SAGE results shown are zonal averages obtained from about 30 measurements in a month. The CCD measurements are from the Pacific only. (The implicit assumption here is that the stratospheric ozone column is longitudinally invariant.) Nonetheless, there is a striking agreement between these two independent measurements. Given these measurements one can then estimate tropospheric O3

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FIGURE 6. The tropical deep convective cloud model proposed by Kley et al. [1996] to explain their observation of near-zero ozone mixing ratio in the tropical transition layer (TTL). These clouds allow a simple yet powerful method of estimating stratospheric O3 overburden that can be subtracted from total column O3 (measured over cloudfree scenes) to study tropospheric column O3.

column from total column O3 measured over clear scenes. Figure 8 compares the times series derived this way with the time series from the balloon instruments measured at two tropical locations. Ziemke et al. [2001] have extended this method to a more general “cloud slicing” method in which clouds at multiple heights in the atmosphere can be used to estimate mean mixing ratio of O3 in the upper and lower troposphere.

Nitrogen Dioxide (NO2) NO2 is an inevitable byproduct of combustion, produced by automobiles and biomass burning; it is also produced by lightning. Though, by itself, NO2 is not considered harmful, it is the primary source of atomic oxygen in the troposphere that leads to ozone production. The mixing ratio of NO2 in the troposphere varies from less than 0.1 parts per billion to more than 10 per billion in highly polluted urban areas. However, since most of the NO2 is vertically localized, the total column NO2, typically measured by satellites, rarely exceeds few times 1016 molecules/cm2, producing less than 1% reduction in TOAR (with respect to the background value) at ~400 nm, where NO2 has maximum absorption. To make matters worse, from the point of view of satellite instruments, highest concentration NO2 usually occurs in areas of less than 100 km2 and remains localized in the planetary boundary layer (PBL) and is thus easily obscured by clouds. Yet,

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FIGURE 7. Twenty-year time series of stratospheric column O3 derived using Pacific deepconvective clouds (CCD method) is compared with SAGE measured column O3. These two independent methods agree remarkably well; both capture the very significant difference in behavior of the O3 time series between 1 °S (upper panel) and 5 °S.

remarkably, it has been possible to detect NO2 from satellite instruments [Richter & Burrows, 2002; Wenig et al., 2004]. Figure 9 shows a remarkable global map produced from data collected by the SCIAMACHY instrument on the ENVISAT satellite, showing the NO2 pollution generated by all the major urban centers of the world. Despite these noteworthy successes in producing such maps, the accuracy and precision of the derived product remain marginal for careful quantitative studies of AQ. From the measurements made so far it is particularly difficult to track the NO2 plumes away from the source regions; such transports are responsible for converting a local air pollution problem into a regional and even a global problem.

Carbon Monoxide (CO) In principle, CO can be measured using the 2.3-micron absorption band where there is still a sufficient amount of sunlight to measure the reflected radiation. However, this band is extremely weak and so far the experience in using this band has been limited. (There are preliminary indications that the SCIAMACHY instrument is able to see CO using this band.). So far, satellite measurements of CO have come primarily from the 4.6-micron thermal emission band of CO,

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FIGURE 8. Twenty-year time series of tropospheric column O3 derived using the two tropical locations. The CCD methods capture the very significant difference in behavior of the trop O3 time series between Nairobi (upper panel) and Natal. These results give confidence that clouds can be used as a powerful tool to “slice” the atmosphere to study tropospheric O3.

which provides information about CO in the middle troposphere. Recent results have come from the MOPITT instrument on the EOS Terra and the AIRS instrument on the EOS Aqua satellite.

Sulfur Dioxide (SO2) There are both natural and anthropogenic sources of SO2. Large amounts of SO2 are routinely emitted by fuming volcanoes, but they are short-lived and cause localized pollution. Volcanic eruptions can inject SO2 into the upper troposphere and stratosphere where they quickly convert to sulfuric acid aerosols that may last a long time, even years, causing perturbation to climate and radiation. These high-altitude injections of SO2 are fairly easy to monitor from satellites. Even the Nimbus-7 TOMS instrument, which was not designed to measure SO2, but was later found to have weak sensitivity to it, has seen SO2 injected into the atmosphere by about one hundred volcanic eruptions during its 13.5-year lifetime [Carn et al., 2003]. The more recent instruments—GOME, SCIAMACHY and OMI—have more than an order of magnitude greater sensitivity to SO2 than TOMS. In particular, the recently launched OMI instrument, which has a smaller pixel size than its predecessor instruments, should be able to monitor the fuming volcanoes.

1. Air-Quality Study from Geostationary/High-Altitude Orbits

Tropospheric NO2 column (1015 molec cm−2) April 2004

KNMI / IASB / ESA

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4.0

6.0

8.0

10.0 20.0

FIGURE 9. April 2004 map of tropospheric NO2 produced from data collected by the SCIAMACHY instrument. (Map courtesy of KNMI, NL.)

The second major source of SO2 is the burning of sulfur-rich coal. Although many industrialized nations have reduced the emission of SO2 considerably in the last few decades by using sulfur scrubbers in their power plants, many countries in the developing world do not do so. In addition, in many countries, burning of sulfur-rich coal in homes and small factories are major contributors of SO2. However, given the very short lifetime of SO2 near the surface, with rare exceptions [Carn et al., 2004], anthropogenically generated SO2 doesn’t usually build up in large enough amounts to be seen easily from satellite instruments. Yet, as Figure 10 illustrates, it has been possible to measure SO2 near the surface from satellites in some parts of the world where the SO2 pollution is most severe. However, to increase the scientific value of such measurements it will be necessary to increase the sensitivity of these measurements by an order of magnitude.

Measurement from High Vantage Point Satellite Orbits With the exception of aerosols, satellite measurements of atmospheric pollutants have been largely confined to the Low-Earth Orbit (LEO), roughly defined as orbits with altitude of less than 1000 km from the surface of the Earth. These orbits have several inherent problems for AQ studies. To measure over a groundswath of, say, 1500 km, from a typical 750-km altitude orbit, requires a camera of field-of-view (FOV) of about 90 ° (assuming one is using a two-dimensional CCD camera, similar to the popular digital cameras of today). In 35-mm camera

34

Pawan K. Bhartia GOME: Maximum SO2 January 2000

VC

SO2

[DU]

48

5.00 44

4.50 4.00

40

3.50 36

3.00 2.50

32

2.00 28

1.50 1.00

104

108

112

116

120

128 124 Longitude

132

136

140

144

FIGURE 10. January 2000 map of SO2 produced from data collected by the GOME instrument. (Map courtesy of Prof. John Burrows, U. of Bremen, Germany.)

terms, this amounts to using a 17.5-mm wide-angle lens. To get decent optical performance from such wide-angle optics, one is forced to use very small apertures (called the f/number), which limits the amounts of photons that can be collected. A typical aperture size for such a camera might be ~2 mm. Another serious problem with LEO instruments is that the satellite moves too fast (~7 km/sec), which severely limits the integration time (“shutter speed”); similar to taking the picture of a nearby object from a fast moving vehicle without blurring the picture. Finally, an LEO instrument (measuring reflected sunlight) with 1500km swath will see any point on the Earth’s surface once every two days, and perhaps every 4 days if cloud free pixels are required. Although the coverage can be improved by increasing the swath, it would be of little value for AQ studies for the pixels at the edges of the swath become much too large and are frequently cloud contaminated. Since AQ changes on a hourly scale, the temporal coverage provided by LEO instruments are deemed inadequate to study AQ. Moving farther away from the Earth brings in several advantages. To start with, let’s first dispose off the myth, held by many, that moving away from the Earth would reduce the signal/noise. Though the number of photons go down by the square of the distance, as users of SLR cameras know well, the number of photons received by a camera is a function of the f/number and not of its distance from the object; the distance effect is canceled exactly by the fact that longer focal length lenses of the same f/number have larger apertures. As we discussed above, the LEO instrument actually require a slower optics (larger f/number) to minimize optical distortion, which is not a problem for measurements made from larger distances. For example, if one were to move 40 times away

1. Air-Quality Study from Geostationary/High-Altitude Orbits

35

from Earth, from 750 km to 30,000 km, one would require a modest 80-mm aperture instead of the 2-mm aperture (forced by the wide angle optics) to get exactly the same swath and S/N. And, since the FOV of the instrument would also decrease by a factor of 40, one could easily design a faster optics to improve S/N. Moreover, by the Kepler’s law, satellites in more distant orbits move slower, so one can use a slower shutter speed, increasing the integration time to achieve even better S/N. Finally, and perhaps most importantly, a point on the Earth remains visible longer the farther one goes away, so one can pick and choose areas one wants to study and then study the area for long periods of time. One could, for example, wait for clouds to move away from an area to see what is underneath or for clouds to come over an area to isolate what is above them. Indeed the cloud slicing becomes a very compelling tool for sounding the troposphere from a high vantage point orbit. The table in Figure 11 shows that there are many tradeoffs in choosing a high vantage point orbit. The primary one is the observation time versus spatial coverage.

Orbit Scenarios LEO

MEO

GEO

L1 1,60,000

Altitude (km)

564

20,187

35,784

No.of orbits/day

15

2

1

Earth’s rel motion (/orbit)

24⬚

−360⬚

0

0 −360⬚

Max scan angle (sat za = 70⬚)

59.7

13.0

8.2

0.2

Swath (km)

2,293

12,662

13,742

15,508

Obs. Time (hr/pixel)- TIR

0.09

7.6

24.0

9.3

Obs. Time (hr/pixel)- UV/VIS/NIR

0.09

5.7

12.0

9.3

Daily Spatial Coverage

global

global

13,742 km

global

Sun-glint obs

all lats

Multi-angle obs

all lats

tropics

No

tropics

1

36

63

2,837

Aperture in cm (200:1 S/N) (5 km pixel, 0.5 sec integ time)

sub-solar lats sub-solar lats

sub-solar lats

FIGURE 11. The table shows the tradeoffs in choosing the orbit for AQ studies from satellites. For a nominal measurement requirement (2000:1 S/N at ~400 nm, 5-km pixel, 0.5sec integration), one requires an instrument of ~1-cm aperture in 564-km orbit. (For reasons discussed in the text, it is difficult to build such an instrument if the required swath is 2293 km, which is needed to get global coverage.) As one moves farther away from Earth to 20,187 km MEO, one continues to get global coverage with increasing observation times per pixel. (Note that the obs time given for UV/VIS/NIR instruments in MEO are mean values; some pixels will get substantially more and some substantially less observation time.) Going to GEO increases the observation time further but reduces the coverage. Going to L1 provides a modest increase in observation time (over MEO), but requires unacceptably large apertures.

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Pawan K. Bhartia

While the geostationary orbits are ideal for AQ studies, for they keep an area perpetually in the field-of-view, they provide measurements over less than onefourth of the Earth’s surface area. By moving to lower orbits one allows the orbit to drift over the entire globe but reduces the temporal coverage. An orbit location that some people find attractive is the 1st Langrange point (L1), which lies at ~1.6 million km from the Earth along the Earth–Sun line. For instruments that work with reflected sunlight this location provides full global coverage in one day, with each sunlit pixel remaining in view for close to 9 hours as the Earth revolves under the satellite. However, as the table shows, one requires unacceptably large aperture from this vantage point to study AQ, perhaps with the exception of aerosols, which require a more modest S/N than the others.

Summary The AQ is controlled primarily by the concentrations of particulates smaller than 2.5 micron and of 4 trace gases that include O3, SO2, CO, and NO2. Satellite instruments, primarily operating from low orbits (< 1000 km altitude), have demonstrated that all five constituents can be measured from space. However, these measurements lack the spatial and temporal resolution deemed necessary for AQ studies. The high vantage point orbits (> 20,000-km altitude) are highly attractive for AQ studies, for instruments in these orbits can provide good temporal and spatial resolution with improved S/N than the LEO instruments. Maximum temporal coverage is provided by instruments in the Geostationary (GEO) orbit, but they can observe less than one-fourth of the globe, thus requiring multiple satellites for global coverage. UV/VIS/NIR instruments, which use reflected sunlight to make measurements, can use cloudy pixels to sense the atmosphere above the clouds, and cloudfree pixels to sense the atmosphere below the clouds, thus providing some altitude information. Thermal IR instruments have the advantage that they can operate 24 hours a day, but they are sensitive to species in the middle troposphere only, and require cloudfree pixels, which severely limits the useful coverage they can provide from LEO. The high vantage points orbits can greatly improve the number of cloudfree observations one can make, improving the utility of these instruments for AQ studies. Combination of reflected-sunlight and thermal emission instruments would provide highly synergistic measurements of AQ from space.

References Carn et al. (2003), Volcanic eruption detection by the Total Ozone Mapping Spectrometer (TOMS) instruments: A 22-year record of sulphur dioxide and ash emissions, Oppenheimer, C., Pyle, D.M. & Barclay, J. (eds) Volcanic Degassing. Geological Society, London, Special Publications, 213, 177–202. 0305-8719/03.

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Carn S.A., A.J. Krueger, N.A. Krotkov, and M.A. Gray (2004), Fire at Iraqi sulfur dioxide plant emits SO2 clouds detected by Earth Probe TOMS. Geophys. Res. Lett., 31, L19105, doi:10.1029/2004GL020719. Fishman, J., C.E., Watson, J.C. Larsen, and J.A. Logan (1990), Distribution of tropospheric ozone determined from satellite data, J. Geophys. Res., 95, 3599–3617. Herman, J.R., P.K. Bhartia, O. Torres, C. Hsu, C. Seftor and E. Celarier [1997], Global distribution of UV-absorbing aerosols from Nimbus7/TOMS data, J. Geophys. Res., 102, 16911–16922. Hsu, N.C., J.R. Herman, P.K. Bhartia, C.J. Seftor, O. Torres, A.M. Thompson, J.F. Gleason, T.F. Eck, and B.N. Holben (1996), Detection of biomass burning smoke from TOMS measurements, Geophys. Res. Lett., 23, 745–748. Hsu, N.C., S. Tsay, M.D. King, and J.R. Herman (2004), Aerosol properties over brightreflecting source regions, IEEE Trans. Geosci. Remote Sens., 42, 557–569. Kley, D. et al. (1996), Observations of near-zero ozone concentrations over the convective Pacific: Effects on air chemistry, Science, 274, 230–232. McPeters, R.D., et al. (1996), Nimbu&stilde;7 Total Ozone Mapping Spectrometer (TOMS) Data Product’s User’s Guide, NASA Reference Publication 1384, National Aeronautics and Space Administration, Washington, DC. Remer, L.A. et al. (2004), The MODIS aerosol algorithm, products and validation, to be published in J. of Atmos Sci, Special issue: Chesapeake Lighthouse and Aircraft Measurements for Satellites (CLAMS) Field Experiment. Richter, A. and J. Burrows (2002), Retrieval of tropospheric NO2 from GOME measurements, Adv. Space Res., 29(11), 1673–1683. Torres, O., P.K. Bhartia, J.R. Herman, Z. Ahmad, and J. Gleason (1998), Derivation of aerosol properties from satellite measurements of backscattered ultraviolet radiation: Theoretical Basis, J. Geophys. Res., 103, 17,099–17,110. Wenig et al. (2004), Retrieval and analysis of stratospheric NO2 from the global ozone monitoring experiment, J. Geophys. Res., 109, D04315, doi:10.1029/2003JD003652. Ziemke, J.R., S. Chandra, and P.K. Bhartia (1998), Two new methods for deriving tropospheric column ozone from TOMS measurements: The assimilated UARS MLS/HALOE and convective-cloud differential techniques, J. Geophys. Res., 103, 22,115–22,127. Ziemke, J.R., S. Chandra, and P.K. Bhartia (2001), “Cloud slicing”: A new technique to derive upper stratospheric ozone from satellite measurements, J. Geophys. Res., 106, 9853–9867.

Chapter 2 Aerosol Forcing and the A-Train CHIP TREPTE

Introduction Aerosols are small particles consisting of solid or liquid material suspended in air, with diameters ranging from about 0.01 to 1 µm. They originate primarily from natural sources such as volcanoes, dust storms, forest fires, vegetation, and sea spray. Anthropogenic sources include burning of fossil fuels and land use changes such as biomass burning and deforestation. Depending upon their size, injection altitude, and proximity to precipitating cloud systems, lifetimes of aerosols in the atmosphere vary considerably from a few hours to as much as 2 weeks and thus, can be transported long distances from their sources. Aerosols can have a significant impact on the environment and climate. They can have a direct radiative forcing effect by scattering and absorbing sunlight. The change in direct radiative forcing due to anthropogenic aerosols alone (also known as direct aerosol climate forcing) is substantial and rivals that due to stable greenhouse gases. For the most part it is opposite in sign to the greenhouse gases, however, the regional nature of aerosol forcing and its variable vertical distribution precludes simple conclusions of a global canceling of greenhouse gas warming. Aerosols can also indirectly influence radiative forcing by modifying the albedo of clouds either through the incorporation of absorbing aerosols, which tend to darken them, or through the seeding of smaller droplets, which tend to brighten them. It has been suggested that aerosols can alter the liquid water content of clouds, thereby lengthening cloud lifetime and the geographical extent of cloudiness, or even change the vertical structure of latent heating and influence atmospheric dynamics. At present, our ability to quantify and predict the impact of global aerosol forcing on climate remains highly uncertain [IPCC, 2001]. The uncertainties arise, in part, because of insufficient knowledge of aerosol variability, composition, optical properties, hygroscopicity, and size distribution. Complex aerosol processes such as transport, transformation, aging, and cloud interactions are also poorly represented in models. More observations from a variety of techniques, platforms, and

38

2. Aerosol Forcing and the A-Train

39

vantage points are needed to improve our characterization of the aerosol system and reduce modeling uncertainties. In the next 12 months, the A-train satellite constellation consisting of the Aqua, CloudSat, CALIPSO, PARASOL, and Aura satellite missions will be fully assembled in orbit and will provide valuable new information to help address some of these uncertainties. For the first time, a three-dimensional view of the aerosol and cloud system over the globe will be available from a combined instrument suite including a lidar, a radar, multispectral imagers and broadband radiometers. These measurements will enable more accurate estimates of direct aerosol forcing than presently available. They will be made above, below, and next to cloud systems; over bright and dark surfaces; and at high and low latitudes. The measurement suite will further permit a better understanding of regional impacts of indirect aerosol forcing and the evolution of aerosol distributions that will help to constrain parameterizations of aerosol processes in climate models. This lecture will provide an overview of the basic factors governing direct and indirect aerosol radiative forcing. It will be followed with a summary of the measurement capabilities for aerosols on the A-Train constellation and on how these might be utilized to better understand aerosol processes and provide more accurate estimates of aerosol forcing.

Direct Aerosol Forcing In the early 1990s it was believed that radiative forcing by tropospheric aerosols was dominated by sulfate droplets such as those found in the stratosphere following large volcanic eruptions. These aerosols absorbed negligible amounts of visible light and, hence, have a single scattering albedo (fraction of light attenuated by scattering to the total amount attenuated by scattering and absorption) close to unity [Charlson et al., 1992]. Over the years, however, greater realization of the importance of different aerosol types and their optical properties revealed that aerosol forcing is more complicated and may, under suitable conditions, cause net heating. The following expression [Haywood and Shine, 1995] is used to identify how different factors influence aerosol radiative forcing: 1 - 1) DF = - DS o T at2 (1 - A c ) ~ b d e (1 - R s ) 2 - 2R s ( ~ o b where D is daylight, So is the solar constant, Tat is the atmospheric transmission above the aerosol layer, Ac is fractional cloud cover, Rs is the surface reflectivity, ω is the single scattering albedo (fraction of light attenuated by scattering to the total amount attenuated by both absorption and scattering), δ is optical depth, and the β is the upscatter fraction. The expression reveals that in clear skies ∆F is governed principally by aerosol abundance through δ, aerosol optical properties through β and ω, and the underlying surface reflectance. Further inspection of the above equation reveals that forcing can shift from net cooling to net warming when

40

Chip Trepte

~<

2R s b (1 - R s ) 2 + 2R s

Typical values of the upscatter fraction β range from .15 to .30. For sulfate aerosols (ω ~ 1), the radiative forcing is negative over almost all surface conditions. However, for aerosols that absorb solar radiation (e.g., containing black carbon) the single scattering coefficient decreases, net warming occurs when the particles overlie gray or bright surfaces. For example, when ω = 0.9 and Rs > 0.3, the atmosphere will warm. This example is simplified to illustrate a few of the complexities involved in modeling direct aerosol forcing. More accurate representation would include factors with possible interdependencies such as hygroscopicity, internal/external mixtures, varying solar angles, etc. Since climate forcing is due only to the anthropogenic component of the aerosol distribution, we need isolate this contribution from the total radiative forcing calculation. This aspect is best performed by chemical analysis with in situ instrument probes. An upper estimate of global climate forcing from satellite observations may be obtained by realizing that anthropogenic aerosols often fall within the fine mode fraction as suggested by a conceptual representation of the aerosol size distribution. Coarse mode particles are usually formed by mechanical actions (e.g., wind-blown dust or sea salt). Fine mode particles often form as a result of combustion or nucleation/condensation. The difficulties of isolating the anthropogenic component and the need for global observations underlines the need for an integrated measurement sensor web that includes ground-base, aircraft, satellite observations fused together by data assimilation techniques [Diner et al., 2004; Anderson et al., 2004].

Indirect Aerosol Forcing Since aerosols are the building blocks of clouds, it is natural to expect that changes in aerosol characteristics can impact the properties of clouds. Twomey [1977] first described how the presence of more cloud condensation nuclei could increase the liquid water content (LWC) of clouds and change their cloud optical properties. To illustrate the relationship, consider the definition of optical depth, 3

d = rh

#Q r e

2

n (r) dr

0

where h is the depth of the cloud, Qe is the extinction efficiency, r is the particle radius, and n(r) is the size distribution. At visible wavelengths Qe ~ 2. If we further consider a narrow droplet-size spectrum with a mean radius, r-, the above expression can be approximated as d = 2r h r 2 N where 3

N=

# n (r) dr 0

2. Aerosol Forcing and the A-Train

41

is the total number concentration of droplets. The total liquid water content of a cloud is defined as 3 LWC = 4r t # r 3 n (r) dr 3 0 After making similar substitutions for the droplet size distribution and mean radius the expression can be approximated by LWC = 4r t r 3 N 3 Relating together the approximated expressions for optical depth and liquid water content yields 2/3 d = 2.4 c LWC hN m t This has the consequence that an increase in cloud condensation nuclei produces an increase in optical thickness. Because of the increased competition for water vapor with the increased aerosol population, cloud droplets may be unable to grow large enough to initiate the collision-coalescence process (the dominate precipitation process for warm clouds). The prevailing thought is that precipitation rates, thus, will decrease and cloud lifetimes lengthen. A number of studies have observed changes in aerosol concentrations, cloud droplet number/size, and cloud reflectance. Many of these studies include both aircraft and satellite observations at near-infrared wavelengths, where cloud reflectance is highly sensitive to droplet size rather than optical thickness [e.g., Nakajima et al., 2001]. Some of the most compelling evidence showing changes in aerosol optical properties by aerosol concentration comes from AVHRR observations of the effects of ship exhaust on clouds [Coakley et al., 1987]. Because of the magnitude and degree of uncertainty with indirect aerosol forcing, more observations and more modeling research is needed. Further investigations from the advantage point of space, hopefully, can continue to provide an improved understanding of the regional and global affects of aerosol-cloud interactions and their impact on climate.

A-Train Aerosol Measurements The A-Train satellite constellation (named for the Aqua and Aura satellites) has an orbit at an altitude of 705 km and inclination of ~98°. Aqua leads the constellation with an equatorial crossing time of approximately 1:30 p.m. and Aura trails the set with a crossing time of ~1:45 p.m. On December 2004, the PARASOL satellite was launched followed by the joint launch of CALIPSO and CloudSat on April 2006. These satellites have been joined the constellation in orbit between the Aqua and Aura platforms. Table 1 summarizes the instruments on the different platforms and products available for aerosol studies. As seen in the table, a large number of near-coincident measurements on aerosol and cloud physical and optical properties, radiative flux, and thermodynamic parameters will be available over the same locations within a few minutes of each other. (Details on the

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Chip Trepte

A-train constellation may be found at the following Web site: http://eospso.gsfc. nasa.gov/eos_homepage/for_educators/educational_publications.php.) Global maps of aerosol optical depth can be derived from measurements by MODIS, PARASOL and OMI. These instruments exploit the spectral sensitivity of the size distribution and composition of aerosols to infer their optical properties. For MODIS, visible and near-IR measurements over dark surfaces are used to retrieve aerosol optical depth [Kaufman et al., 2002] and provide discrimination between accumulation and coarse size modes [Remer et al., 2002]. Over bright surfaces, less spectral information is available because of the lack of spectral contrast and the complications of large variability in surface bidirectional reflectance. An example map of aerosol optical depth derived from MODIS is shown in Figure 1 for the months of January and July 2003. For the month of January, most oceanic regions have δ < 0.2. However, near the western coast of Africa and over central Africa values peak around 0.8. High values are also present over eastern China. For the month of July, much higher aerosol loading is present over the globe with peaks in regions near or downwind of the major deserts and large areas of industrial activity. The white regions in both figures are data void areas (note that they occur over high albedo regions). It should be noted that application of the technique is permitted only for cloud free pixels. The data shown represents the cumulative average for available scenes for the month. These were obtained from a catalogue of aerosol data products from the MODIS Online Visualization and Analysis System at (http://lake.nascom.nasa.gov/www/ online_analysis/movas). Aerosol retrievals using POLDER observations on PARASOL are expected to follow the technique demonstrated Bellouin et al. [2003]. The approach uses multiangle images at 9 different wavelengths (443, 490, 565, 670, 763, 765, 865, 910, and 1020 nm) with channels at 490, 670, and 865 nm sensitive to polarized signals to derive optical depth and information on aerosol size distribution (Angstrom coefficient) over the ocean. Over land, the technique provides an index that is proportional to the product of optical depth and the Angstrom coefficient. The OMI aerosol retrieval uses near-UV spectral channels instead of those used by MODIS in the mid-visible region. The advantage in this approach is that most land surfaces are dark in the near UV and the interaction between aerosol and Raleigh scattering offers sensitivity to aerosol absorption [Torres et al., 2002]. This feature permits OMI to provide insight on the bulk distribution of single scattering albedo over the globe. The addition of CALIPSO mission to the A-train constellation greatly complements the other aerosol observations by providing information on the vertical structure of aerosol and cloud layers [Winker et al., 2003]. CALIPSO consists of a polarization-sensitive lidar operating at 1064 and 532 nm and passive imagers operating in the visible and infrared spectral regions. It has a designed optical depth sensitivity of at least 0.005 at 532 nm. With this information, aerosol features can be unambiguously identified above and between clouds and over bright surfaces. The technique is especially sensitive at low optical depths and promises

2. Aerosol Forcing and the A-Train

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[unitless] (January 2003) Aerosol Optical Thickness 90N

60N

30N

EQ

30S

60S

90S 180

120W

0 GrADS: CCLA/IGES

0.1

60W 0.2

0.3

0 0.4

0.5

60E 0.6

0.7

120E 0.8

0.9

180 1 2004−10−22−09:50

[unitless] (July 2003) Aerosol Optical Thickness 90N

60N

30N

EQ

30S

60S

90S 180

120W

0 GrADS: CCLA/IGES

0.1

60W

0.2

0.3

0

0.4

0.5

60E

0.6

0.7

120E

0.8

0.9

180

1 2004−10−22−09:39

FIGURE 1. Aerosol optical depth obtained from the MODIS satellite experiment for the months of January and July 2003.

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Chip Trepte

to improve our quantitative estimate of global aerosol forcing. This is illustrated in Figure 2, which shows the estimated uncertainty derived from CALIPSO compared with retrievals from MODIS. At low optical depths (< ~.2), lidar retrievals are superior. Their accuracy degrades at higher optical depths because of greater weighting in the uncertainty of the extinction-to-backscatter coefficient that is used in the aerosol retrieval algorithm. This parameter is dependent on composition and the size distribution. CALIPSO’s measurements of aerosol layers will also be valuable in evaluating model transport processes. It is widely recognized that chemical-transport models still have difficulty in reproducing observed vertical distributions of aerosols. The ability to distinguish aerosol layers within or above the boundary layer will help to identify transport pathways and provide insight on their region of origin with the use of backtrajectory analysis and tracer observations. An example of a set of lidar observations is shown in Figure 3 from the LITE experiment, which flew on the Space Shuttle in 1994. In this figure, the vertical structure of clouds and aerosols are clearly evident. The yellow lines show backtrajectories and highlight the nature of differential transport processes in the atmosphere. Another area where the CALIPSO data will provide new information is in the use of the ratio of 532/1064-nm backscatter measurements. According to Mie 1.0 CALIPSO (Ext-to-Backscatter error - 3096) MODIS, Land (Kaufman, 1997) MODIS, Ocean (Tanne, 1997)

Uncertainly, ∆τ/τ

0.8

0.6

0.4

0.2 Globel-mean derived by MODIS 0.0 0.0

0.2

0.4

0.6

0.8

1.0

Aerosol Optical Depth, τ

FIGURE 2. Uncertainty estimate for retrieved aerosol optical depth. The CALIPSO algorithm assumes an uncertainty of 30% in the extinction-to-backscatter coefficient. The two different error estimates from MODIS reflect the use of different retrieval techniques for either land or oceanic surfaces.

2. Aerosol Forcing and the A-Train

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FIGURE 3. Example cross-section of lidar backscatter observations obtained from the LITE experiment over North America. Aerosol layers are denoted by warm colors. Clouds are indicated by white features. The yellow lines are backtrajectories.

theory, ratios approaching unity suggest the presence of larger particles, while those at lower values suggest the presence of smaller particles. This information coupled with absolute backscatter or extinction values can point to the likelihood of cloud, coarse, or fine mode particles. From this data, clouds and aerosols can be separated with confidence as a function of altitude. This capability may prove valuable in tuning of cloud mask algorithms needed for aerosol and cloud retrievals from MODIS and POLDER. For understanding interactions between aerosols and clouds, CloudSat will provide key information on the properties of optically thick clouds, their ice and liquid water content, and associated rates of precipitation with its 94-GHz radar [Stephens et al., 2002]. Reflectivity measurements will be obtained from the surface to 30 km with 500-m vertical resolution. Because cloud properties can change rapidly, the CloudSat platform will be placed in orbit ~2 minutes behind Aqua, with CALIPSO trailing CloudSat by ~15 sec. In addition, control of CloudSat’s cross-track motion will be within ±1 km of the CALIPSO ground track to ensure the most lidar/radar observations are co-aligned. Concurrent observations of thermodynamic state parameters (e.g., temperature, relative humidity) from instruments on Aqua will provide the ability to examine interrelationships between aerosol optical depth, single scattering albedo, and relative humidity. Importantly, concurrent A-train aerosol observations with radiative flux measurements from CERES will enable more accurate

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Chip Trepte

TABLE 1. A-Train Instrument Characteristics and Related Products Spacecraft Aqua

Instruments MODIS

Characteristics 36-channel visible radiometer 2300-km swath Variable pixel resolution .25 to 1 km AIRS/AMSU-A IR and microwave sounders, 1650-km swath IR pixel resolution ~10 km AMSR-E 6-channel microwave radiometer 1445-km swath CERES

HIRDLS

Broadband and spectral radiances converted to fluxes Pixel resolution ~20 km 500-m vertical range gates from surface to 30 km Pixel resolution ~1.4 km 532- and 1064-nm channels with depolarization Field of view 70 m at surface with 333-m separation between samples 3-channel IR radiometer Visible wide field of view camera 64-km swath IR pixel resolution 1 km WFC pixel resolution 125 m within 2.5 km of nadir and 1 km beyond 9-channel polarimeter with channels in visible and near IR. 400-km swath Pixel resolution 5 km IR limb sounder

MLS

Microwave limb sounder

TES

IR imaging spectrometer Pixel resolution .5 × 5 km UV grating spectrometer

CloudSat

94-GHz radar

CALIPSO

Lidar

IIR/WFC

PARASOL POLDER

Aura

OMI

Aerosol/Cloud Products Land, ocean, and atmospheric products; include cloud and aerosol optical depths and particle size information Temperature and moisture profiles in clear atmosphere Liquid water path, column water vapor, liquid precipitation Top of atmosphere radiation budget Time mean and instantaneous fluxes Cloud profile information, liquid and ice water content, precipitation Aerosol and cloud profile for optically thin clouds Aerosol extinction, optical depth, and typing Cirrus cloud properties

Cloud and fine mode aerosol optical depth and particle sizes

Trace gases and stratospheric aerosol Trace gases, ice content of upper tropospheric cloud Trace gases Ozone, aerosol optical depth and single scattering albedo

Pixel resolution 13 × 24 km

observationally based estimates of aerosol forcing over the globe. This method has already been demonstrated by Haywood et al. [1999] and Christopher and Zhang [2002] using MODIS/CERES data. The added information from CALIPSO and CloudSat should improve estimates near and in cloudy regions and regions with bright surfaces.

2. Aerosol Forcing and the A-Train

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Acknowledgements Some of the data used in this lecture were acquired as part of the NASA’s Earth Science Enterprise. The MODIS aerosol retrievals were provided by MODIS Science Teams and data processed by the MODIS Adaptive Processing System (MODAPS). These are archived and distributed by and Goddard Distributed Active Archive Center (DAAC).

References Bellouin, N., O. Boucer, D. Tanre and O. Dubovik, Aerosol absorption over the clear-sky oceans deduced from POLDER-1 and AERONET observations, Geophys. Res. Lett., 30(14), 1748, doi:10.1029/2003GL017121, 2003. Charlson, R.J., S.E. Schwartz, J.M. Hales, R.D. Cess, J.A. Coakley, J.E. Hansen, and D.F. Hofmann, Climate forcing by anthropogenic aerosols, Science, 255, 423–430, 1992. Coakley, J.A., R.L. Bernstein, and P.A. Durkee, Effect of shipstack effluents on cloud reflectivity, Science, 237, 1020–1022, 1987. Husar, R., J.M. Prospero, and L.L. Stowe, Characterization of tropospheric aerosols over the oceans with the NOAA advanced very high resolution radiometer optical thickness operational product, J. Geophys. Res., 102, 16889–16909, 1997. Intergovernmental Panel on Climate Change, Aerosols, their direct and indirect effects, ed., J. Penner et al., Climate Change 2001, The Sciencitific Basis. Working Group I to the Third Assessment Report of the IPCC, Camridge University Press, 2001. Kaufman, Y.J., and T. Nakajima, Effect of Amazon smoke on cloud microphysics and albedo – Analysis from satellite imagery, J. Appl. Meteor., 32, 729–744, 1993. Nakajima, T., A. Kawamoto, and J.E. Penner, A possible correlation between satellitederived cloud and aerosol microphysical parameters, Geophys. Res. Lett., 2001. Remer, L.A., D. Tanre, Y.J. Kaufman, C. Ichoku, S. Mattoo, R. Levy, D.A. Chu, B. Holben, O. Dubovik, A Smirnov, J. Martins, R. Li, and Z. Ahman, Validation of MODIS aerosol retrieval over ocean, Geophys. Res. Lett., 29, 10.1029/2001GL013204, 2002. Stephens, G.L., D.G., Vane, R.J. Boain, G.G. Mace, K. Sassen, Z. Wang, A. Illingworth, E O’Conner, W. Rossow, S. Durden, S. Miller, R. Austin, A. Benedetti, C. Mitrescu, The CloudSat mission and the a-train; a new dimension of space-based observations of clouds and precipitation, Bull. Amer. Meteor., 83, 12, 1771–1762, 2002. Torres, O., R. Decae, and P. Veefkind, OMI Aerosol Retrieval Algorithm. In OMI Algorithms Theoretical Basis Document, Volume III, Clouds, Aersols, and Surface UV Irradiance, P. Stammes (ed.), NASA-KNMI, 2002. Twomey, S.A., The influence of pollution on the shortware albedo of clouds, J. Atmos. Sci., 34, 1114–1152, 1977. Winker, D.M., J. Pelon, and M.P. McCormick, The CALIPSO mission: Spaceborne lidar for observation of aerosols and clouds, Proc. SPIE, 4893, 1–11, 2003.

Chapter 3 Total Ozone from Backscattered Ultraviolet Measurements PAWAN K. BHARTIA

Introduction The outline of this lecture is as follows. First, we look at how total O3 is derived by looking at the sun from ground-based instruments. This technique was developed by Prof. Dobson in the UK more than half a century ago; still considered the “gold standard” against which all other techniques—both satellite and groundbased—are judged. We discuss several ways of retrieving total O3 from direct sun measurements. Understanding the pros and cons of these methods help us understand the differences between the various methods of producing total O3 from satellite. Finally, we discuss the TOMS total O3 algorithm and show key results produced from this instrument from a record that now spans more than a quarter century.

Total Ozone from the Measurement of Transmitted Solar Flux Referring to Figure 1, the basic quantity one attempts to measure is the atmospheric transmittance T(λ), at wavelength λ, which is defined as the ratio of solar irradiance at the ground F(λ), and the solar irradiance at the top of the atmosphere, F0(λ). From the Lambert–Beer law, T(λ) can be expressed in terms of T0(λ), the transmittance in absence of ozone absorption, multiplied by an exponential term that represents the ozone absorption. The variable m in the exponent is called the air mass factor (AMF), which can be written approximately as secθ0 (by ignoring the sphericity of the Earth’s atmosphere), where θ0 is the solar zenith angle (SZA). The second term in the exponent is the O3 optical depth (τ), which can be written as the product of the ozone absorption cross-section (σO3) and total number of O3 molecules per square area, NO3. σ is usually expressed in cm2, and N in molecules/cm2. However, for total ozone it is customary to use the Dobson Unit (1 DU = 2.67 1016 molecules/cm2), which is named after the British scientist who first measured ozone using the 48

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FIGURE 1. In the direct-sun technique one measures the solar radiation transmitted through the atmosphere, from which the total number of O3 molecules per unit area, N, can be derived using the formula shown.

direct-sun technique [Dobson, 1931]. The amount of ozone in the Earth’s atmosphere However, the large cubic polynomial term in the ozone cross-section complicates the analysis by varying from less than 100 DU inside the Antarctica polar vortex (O3 hole) and can occasionally reach 600 DU in the Feb/March months in the northern hemisphere. Typical values are around 300 DU. In Figure 1, the red band shows a simplified representation of the O3 layer. It is assumed that the O3 molecules are concentrated in a thin layer and are distributed homogeneously throughout the layer. The Lambert–Beer formula, as written, is valid only under this assumption. In the Earth’s atmosphere, ozone is distributed inhomogeneously over a broad range of altitudes. We will consider the impact of this later in this lecture. To estimate NO3, we need a method of estimating T0. At ultraviolet wavelengths where O3 absorption is weak (340–400 nm) Rayleigh scattering by the molecular atmosphere is the principal source of attenuation of the solar radiation, followed by aerosol attenuation, both of which cause T0 to vary smoothly with wavelength. Therefore, if there are no other absorbers besides O3, one can assume that this behavior continues at shorter wavelengths at which O3 is measured (300–340 nm) from solar data. This is the basic assumption made in the Differential Optical Absorption Spectroscopy (DOAS) technique (Figure 2), in which one assumes

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O3 Abs. Cross-section (10−10 cm2)

that the variation of lnT0 with λ can be accurately represented by a 3rd-order polynomial. As shown in the attached figure, with this assumption, one can, in principle, estimate N by linear least-squares regression. However, if one examines the spectral dependence of ozone absorption cross-section (Figure 2 inset graph) one finds that it contains a large 3rd-order polynomial term of its own. So, in the actual fit one estimates N from the residual absorption ∆σ(λ), rather than the total absorption σ(λ). Note that the wiggles in ∆σ(λ) are smaller than 10−20; so, for typical values of Nsecθ0 (~1019 molecules/cm2), one must fit the spectrum to few parts in 104 to achieve 1% accuracy in measuring N. It is very difficult to measure T(λ) to such accuracy even with modern instruments, and was impossible to do so 70 years ago when such measurements got started. However, if one can somehow take advantage of the much larger total absorption, then instrument accuracy/precision requirements can be greatly simplified. The most general method of doing this is to do a two-step retrieval. In the first step, the coefficients of the 3rd-order polynomial representing lnT0(λ) can be derived from measurements taken at wavelengths where the O3 absorption is negligible (340–400 nm); then, in the second step, these coefficients can be used to remove the polynomial from lnT prior to fitting the measurements in the 310–340 nm wavelength range, thus allowing one to estimate NO3 using the total absorption. Although the exact

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FIGURE 2. In the DOAS technique, one estimates N by linear least-squares regression assuming that the attenuation of the solar radiation in absence of ozone can be expressed by a cubic polynomial.

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O3 Abs. Cross-section (10−10 cm2)

method just described hasn’t been widely used so far, the technique used to estimate NO3 from Dobson and Brewer spectrophotometer instruments uses a simplified version of this concept. In the Dobson–Brewer technique, one first reduces the order of the lnT0 polynomial by explicitly accounting for the strong attenuation of the solar ultraviolet radiation by Rayleigh scattering. The Rayleigh optical depth τR (~1 at 310 nm), is proportional to the surface pressure and is known to better than 1%, so it is easy to estimate [Bates, 1984]. The remaining component of T0 is Mie scattering by aerosols Tm. By assuming that lnTm varies linearly with wavelength, one can derive the coefficients of the linear polynomial using a pair of wavelengths at which the O3 absorption is weak; NO3 can then be estimated using a 3rd wavelength. However, for historical reasons, the Dobson–Brewer technique uses two wavelength pairs. Figure 3 shows graphically how the technique works. One notes that the differential absorption one uses in this technique is about 30 times larger than the differential absorption of the DOAS technique; requiring ~1% measurement accuracy to measure NO3 to 1%, greatly simplifying the instrument design and measurement accuracy/precision requirements. Before we end this section, we briefly discuss the impact of the approximations made in Figure 1. As shown in Figure 4, the curvature of the Earth’s atmosphere causes the airmass factor of an atmospheric layer to change with altitude, and the

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temperature dependence of the O3 cross-section causes the O3 absorption optical depth of a layer to do the same. However, one can preserve the Lambert–Beer formula by substituting the average values of m and σO3 in the formula, as shown in Figure 4. To calculate these averages one must have some a priori knowledge of the shape the O3 profile. Since most of the O3 in the Earth’s atmosphere is located between 15–35 km, the a priori dependence is rather small, except at very large solar zenith angles (>80°). Although, for instrumental reasons, Dobson–Brewer measurements are rarely taken above 70° SZA, other ground-based techniques (using zenith sky radiances) can go up to 90° SZA, and direct-sun measurements from high altitude aircraft have been taken up to 92° SZA. Retrieval of NO3 from these measurements requires good a priori knowledge of the shape of the O3 profile. This discussion is designed to highlight the fact that almost all remote sensing techniques require some sort of a priori information to do retrieval. Consequently, the accuracy and precision of the retrieved quantities is determined not just by the accuracy and precision of the basic measurements, but also of the a priori information. This is an important difference between laboratory and in situ techniques and remote sensing techniques. When one studies long-term trends from remote sensing techniques one must consider how long-term changes in the assumed profiles may have impacted the derived trends. It is well known that the shape of the O3 profile has changed over the last 2 decades due to depletion of the stratospheric O3 and increase of the tropospheric O3. This change can affect the trends derived from the remote sensing data and must be considered in the analysis.

FIGURE 4. The curvature of the Earth’s atmosphere and temperature dependence of O3 absorption cross-section causes m and tO3 to become dependent on O3 profile. Retrieval of NO3 requires a priori assumptions about the shape of the O3 profile.

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Total O3 from Satellite Ultraviolet instruments The basic measurement one makes from satellites is the reflectance at the top of the atmosphere, ρ. The definition of ρ given in Figure 5 is such that a perfectly reflecting Lambertian sphere would produce a ρ of unity at all solar zenith angles. Since Earth is not a Lambertian sphere, ρ varies with solar zenith angle, as well as with viewing geometry (satellite zenith and azimuth angles). Although, as shown in Figure 5, the scattering, reflection, and absorption processes that determine ρ are quite complicated, it turns out that, with some approximation, one can write a formula for ρ that looks very similar to the one for the direct-sun transmittance (Figure 1). This leads to retrieval techniques that are conceptually similar to the DS retrieval techniques, though, as we shall see, there are many differences in detail. Figure 6 describes the application of the DOAS technique to derive NO3 from the measurement of satellite reflectances. One immediately notes a new term in lnρ0. This term is introduced by inelastic scattering by the molecular atmosphere (Rotational Raman scattering), in which the wavelength of the incoming and the scattered radiation changes by a small amount (~1 nm). Despite the fact that the wavelength shift is small and only few percent of the scattering is inelastic, this effect (also called the Ring effect) introduces spectral structures in lnρ0 that must be corrected to derive NO3 accurately. In the early versions of the DOAS algorithms, a Ring spectrum was assumed and its magnitude was derived by adding a term in the DOAS fit. More advanced algorithms [Valks et al., 2004]

FIGURE 5. In the 310–340-nm wavelength range, top-of-the-atmosphere reflectance measured by satellite is approximately related to NO3 by a relationship similar to that for directsun transmittance.

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FIGURE 6. Application of the DOAS algorithm to satellite measured reflectances is conceptually similar to that for DS, but requires additional terms, like the Ring effect and correction for some nonlinear effects.

account for the Ring effect caused by the O3 absorption lines themselves (called telluric Ring) to improve accuracy. These algorithms also account for several other effects, some of which are listed in Figure 6, in an approximate way. It is worth mentioning two other more advanced algorithm options, which are described in Figure 7. In the advanced DOAS algorithm, listed as Option 1, the nonlinear contributions to the reflectance from molecular and Mie scattering, surface reflectance, Ring effect, etc., are explicitly calculated using a radiative transfer model, and an a priori ozone profile. The retrieval algorithm then solves for change in NO3 (from the a priori) by a linear least-squares fit on the difference between the measured and calculated reflectances. This algorithm works best for wavelengths longer than 320 nm [Coldewey-Egbers et al., 2004, Yang et al., 2004]. In option 2, one does maximum likelihood estimation to derive an O3 profile using shorter wavelengths, NO3 is then just the integrated amount in the retrieved O3 profile [Bhartia et al., 2004]. The latter algorithm has the advantage that it is less dependent on on a priori O3 profile, particularly at large solar zenith angles.

TOMS Total O3 Algorithm The TOMS total O3 algorithm is a simplified version of the Option 1 algorithm described in Figure 7. The simplification is necessary because like Dobson and Brewer the TOMS instrument has limited number of wavelengths (6), which precludes the application of the full spectral-fitting algorithms described above.

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FIGURE 7. Two advanced spectral fitting algorithms recently developed for processing GOME and SBUV data are briefly described.

Yet, with 25 years of operational experience, the TOMS algorithm has become quite mature and has achieved accuracies comparable to the best-run Dobson–Brewer stations. The concepts developed for the TOMS algorithm have now been adopted by the DOAS and the more advanced spectral-fitting algorithms. This section provides an overview of these concepts. Figure 8 shows the cover of a 39-year-old paper by Dave and Mateer [1967] that first described two key ideas for deriving total ozone from satellite data. With some modifications these ideas were adopted to process early SBUV data [Klenk et al., 1982]. First of these important ideas is the recognition that latitude and total ozone are by far the best predictors of the O3 profile at any given time and location on the Earth [Wellemeyer et al., 1997]; hence they used total ozone itself to select from a set of a priori profiles, starting from a first guess total ozone and then refining it iteratively. Although month has also been added as a predictive variable for profiles in the more recent TOMS algorithm (version 8), as well as in one implementation of the DOAS algorithm [Wellemeyer et al., 2004], concept of using total O3 as a predictive variable for O3 profile was and still remains a powerful concept. Another important idea advanced by the Dave and Mateer paper was to treat Earth’s surfaces and clouds as opaque Lambertian reflectors and to derive their reflectivity using measurements at a weak ozone absorbing wavelength. Dave [1978] later called the reflectivity derived using this assumption Lambertequivalent reflectivity (LER) in recognition of the fact that terrestrial and ocean surfaces are usually non-Lambertian (i.e., their reflectivity varies with solar zenith angle and viewing geometry), and clouds are usually neither opaque nor

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FIGURE 8. First paper describing the technique for measuring total O3 from satellites. TOMS uses a modified version this technique.

FIGURE 9. TOMS V8 algorithm uses 1520 a priori profiles that vary with latitude, total O3, and month. Shown are some examples of these a priori ozone profiles. The altitude at which the ozone partial pressure peaks drops with latitude and total O3.

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Lambertian reflectors. However, they recognized that for the retrieval of total O3 one uses contrast in (top-of-the-atmosphere) reflectances (TOAR) with wavelength caused by O3, rather than the absolute reflectance; therefore, the key issue is not what the LER represents but how does it vary with wavelength. They considered this issue so important that the original TOMS instrument (launched in Oct. 1978 on the Nimbus-7 satellite) was built with 4 wavelengths (331.2, 340, 360, and 380 nm) from which LERs could be derived and compared. By contrast, the instrument had only two O3 sensitive wavelengths (312.5 and 317.5 nm)! In more recent versions of the TOMS algorithm, as well as in most versions of the DOAS algorithms, the LER concept has been modified to handle mixed cloudy scenes better by treating such scenes as mixtures of two Lambert surfaces at two different altitudes; hence called the mixed LER model (MLER). Using Mie scattering theory, Ahmad et al. [2004] show that the MLER model also handles thin clouds better than does the LER model (except over very bright snow-covered scenes). In the basic TOMS version 8 algorithm [Wellemeyer et al., 2004] a weakly ozone-absorbing wavelength (mostly 331.2 nm, but 360 nm at very large SZAs) is used to derive either the LER (when the atmosphere is deemed clear, fully cloudy, or when the ground is covered with snow or ice) or the cloud fraction (fc). (One starts with a nominal total O3 amount to correct for the weak ozone absorption at 331.2 and 360 nm, which is then refined as a better estimate of total ozone becomes available.) Then, assuming that LER/fc is wavelength independent, one can calculate the TOAR (ρ) at any wavelength using an accurate radiative transfer program [Dave, 1964], an a priori ozone profile, and a temperature profile. Total ozone can then be obtained by solving the “Option 1” formula in Figure 7 iteratively, starting with a first guess. However, given limited number of O3 sensitive wavelengths, the TOMS algorithm solves this equation without the polynomial term on the right side of equation. Thus, total O3 is derived from the last term only, using just a single wavelength (mostly 317.5 nm, 331.2 nm at large SZAs). Based on the previous description of the DS technique, it can be seen that this method has the same advantage over the DOAS methods as does the Dobson–Brewer technique; both use the much larger difference in total absorption (between 317.5 and 331.2, or between 331.2 and 360 nm at large SZAs) rather than the residual absorption used in the DOAS methods; this simplifies the instrument design—requiring as few as three wavelengths to retrieve total O3. However, without the polynomial term, the TOMS algorithm is more sensitive to wavelength dependence of LER and instrument radiometric drifts. We discuss these issues below. Wavelength Dependence of LER: Note that the basic TOMS algorithm does not do an explicit correction for aerosols. The rationale for this was provided by Dave [1978]. By explicit Mie scattering calculation, Dave showed that many common types of (non-UV-absorbing) aerosols simply increase the LER of the surface without modifying its spectral behavior, hence they produce no error in the derivation of O3. However, in the same paper Dave showed that aerosols that have strong UV absorption would cause the LER to decrease with wavelength,

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thus such aerosols would introduce error in deriving total O3. Unfortunately, because of interference from subpixel clouds this effect was not discovered in the TOMS data until 1992, 14 years after launch of the first TOMS instrument, when the LER model was replaced with the MLER model, which removed the cloud interference. It is now well established that smoke, desert dust, and volcanic ash strongly absorb UV radiation, and they do produce the effect Dave had predicted. Indeed this observation has led to a unique new way of detecting such aerosols from UV satellite data. The so-called aerosol index, which is derived from the difference in LER between two TOMS wavelengths, is now popularly used for tracking such aerosols over the globe [Hsu et al., 1996; Herman et al., 1997]. (Figure 10 shows an example of how one can track smoke using the TOMS AI) Since these aerosols do produce error in the TOMS algorithm, a correction to the basic TOMS total O3 is applied using the Aerosol Index (AI) as a predictor. (See Torres & Bhartia, 1999 for an earlier version of this correction procedure. The functional form of the correction is somewhat more complex in the TOMS version 8 algorithm.) It is worth mentioning that besides UV-absorbing aerosols highly nonLambertian surfaces, such as the water surface viewed in the geometrical reflection direction (glint), can also cause LER to vary with wavelength. This happens because the LER of such surfaces is very different for direct and diffuse light; from satellite one measures the mixture of the two LERs weighted by the ratio of direct to diffuse radiation. Since this ratio varies with wavelength, so does the

FIGURE 10. This false-color map of TOMS aerosol index (AI) shows smoke from fires in the western US. AI is produced from the difference in LER between two TOMS wavelengths. The color of the land in the background is a visible RGB image from another sensor. Note the absence of clouds in the picture; the TOMS AI algorithm is not sensitive to clouds.

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LER. When the surface of the water is viewed within ~20° of the glint direction, the LER of water surfaces decreases sharply with wavelength, as the direct to diffuse ratio decreases. The magnitude of this effect varies strongly on the surface wind speed, aerosols, and clouds. The AI also allows one to measure the strength of the glint and to correct for its effect on TOMS O3. Radiometric Drift: The radiometric drift (change in instrument sensitivity to incoming radiation) of the instrument can affect the TOMS algorithm in several ways. First, it can cause the derived LER to drift, second it can cause the TOARs (ρ) at the O3 wavelength (317.5 nm) to drift, and third it can cause the AI to drift. All these effects will cause the estimated O3 to drift. Over the years robust techniques have been developed to monitor these drifts. For example, the drift in LER can be monitored by looking at the LER of Greenland and/or Antarctica, both of which are quite stable (though their LERs varies with SZA). The drift in AI can be monitored by ensuring that the AI of bright clouds remains zero with time. Monitoring the drift of the TOAR at 317.5 nm is more difficult. Though several internal techniques have been developed over the years to check the drift in the 317.5 nm TOAR [Herman et al., 1991], ultimately, comparisons with Dobson–Brewer instruments are necessary to maintain data quality over the longterm. This is one area in which the DOAS algorithms may prove superior to the TOMS algorithm.

Key Results from TOMS Figure 11 shows that the first TOMS instrument lasted about 13.5 years [McPeters et al., 1996]. It was followed by two other TOMS instruments. EP/TOMS is currently operating, though it suffers from significant calibration problems that have degraded the data quality (Another TOMS instrument, not shown in Figure 10, was launched on the Japanese ADEOS satellite in 1996, around the same time that EP/TOMS was launched. Since the satellite stopped working after about 9 months, data from this instrument have not been studied extensively.) Figures 12 and 13 show what is generally considered to be the two most important results derived from the TOMS data. Figure 12 shows the long-term changes in the area weighted total ozone between 60S–60N. By combining the data from the both hemispheres, one removes most of the out-of-phase seasonal variation in total O3; some variation that remains is caused by the fact that the amplitude of the seasonal cycle is stronger in the NH than in the SH. Long-term decrease in global total O3 is clearly scene in this data. Although some of this decrease could be due to changes in the meridional transport from tropics to higher latitudes, the bulk of this decrease is believed to be due to increase in the halogen loading of the stratosphere, which causes ozone depletion. The dotted line is a prediction using a two-dimensional stratospheric chemistry model developed at NASA Goddard Space Flight Center. The inter-annual oscillations seen in the model are caused by the effects of solar cycle and changes in aerosol

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TOMS / Ozone Mapper Missions Nimbus 7 TOMS

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FIGURE 12. Long-term change in global (60S–60N) total O3 monitored by TOMS. The dotted line is a model prediction.

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FIGURE 13. Evolution of the area of the Antarctica O3 hole as seen by TOMS.

loading after the eruption of two major volcanoes in the last quarter century: El Chichon in Mexico and Mt. Pinatubo in the Philippines. Figure 13 shows the evolution of the area of the Antarctica ozone hole as seen by TOMS. Though the area varies substantially from year-to-year due to variations in the size and strength of the polar vortex, the long-term increase in area is due to ozone depletion caused by man-made chemicals.

Summary The first paper describing a method for estimating total column O3 from satellite backscatter measurements was published 39 years ago. With advances in computer power, better understanding of atmospheric radiative transfer, and improved instrument design, it is now possible to derive total O3 from satellites with better than 2% accuracy, and ~1% precision over conditions ranging from clear to fully cloudy, up to 85° solar zenith angle, and in presence of massive aerosol loading. There are three basic types of algorithms currently in use. They are briefly described below. Discrete Wavelength Algorithm: This algorithm, derived from Dave and Mateer [1967], has been used to process a quarter-century of TOMS data. The basic algorithm uses only two wavelengths, but 2–3 additional wavelengths are necessary to correct for errors due to aerosols, sea glint, and volcanic SO2. With

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proper selection of wavelengths the effect of atmospheric temperature can be minimized. This algorithm has the advantage that it can be applied to simple instruments that have few spectral channels with relatively coarse spectral resolution (1–2 nm). However, the algorithm is sensitive to instrument drift, which must be monitored by using external reference instruments. Spectral-Fitting Algorithms: These algorithms offer reduced sensitivity to instrument drift but require instruments with narrow bandpass (<0.5 nm), and multiple spectral samples per bandpass, to minimize systematic errors. There are two types of spectral-fitting algorithms. Those that use narrow spectral windows (< 10 nm) are based on the DOAS technique. More advanced algorithms using a broader window (15–20 nm) are being developed to achieve higher accuracy/precision. Maximum-Likelihood Estimation Algorithm: These algorithms derive ozone profile using a broad set of wavelengths (280–320 nm) from which total O3 is obtained by vertical integration. They offer reduced sensitivity to a priori O3 profiles, which becomes a significant error source for the previous two algorithms when solar zenith angle exceeds 80°. However, these algorithms require instruments that can measure weak UV signals down to 280 nm with minimum of scattered/stray light contamination from longer wavelengths.

References Ahmad, Z., P.K. Bhartia, N. Krotkov (2004), Spectral properties of backscattered UV radiation in cloudy atmospheres, J. Geophys. Res., 109, D01201, doi:10.1029/2003JD003395. Bates, D.R. (1984), Rayleigh scattering by air, Planet. Sp. Sci., 32, 785–790. Bhartia, P.K., C.G.Wellemeyer, S.L. Taylor, N. Nath and A. Gopalan (2004), Solar Backscatter Ultraviolet (SBUV) Version 8 Profile Algorithm, Proceeding Quadrennial Ozone Symposium, 1–8 June 2004, Kos, Greece. Dave, J.V. (1964), Meaning of successive iteration of the auxiliary equation of radiative transfer, Astrophys. J., 140, 1292–1303. Dave, J.V. and C.L. Mateer (1967), A preliminary study on the possibility of estimating total atmospheric ozone from satellite measurements, J. Atmos. Sci., 24, 414–427. Dave, J.V. (1978), Effect of aerosols on the estimation of total ozone in an atmospheric column from the measurement of its ultraviolet radiance, J. Atmos. Sci., 35, 899–911. Dobson, G.M.B. (1931), A photo-electric spectrometer for measuring the amount of atmospheric ozone, Proc. Phys. Soc., 324–339. M. Coldewey-Egbers, M. Weber, L. N. Lamsal, R. de Beek, M. Buchwitz, J., P. Burrows (2004), Total ozone retrieval from GOME UV spectral data using the weighting function DOAS approach, Atmos. Chem. Phys. Discuss., 4, 4915–4944. Herman, J.R., R. Hudson, R. McPeters, R. Stolarski, Z. Ahmad, X.Y. Gu, S. Taylor, and C. Wellemeyer (1991), A new self-calibration method applied to TOMS and SBUV backscattered ultraviolet data to determine long-term global ozone changes, J. Geophys. Res., 96, pp. 7531–7545. Herman, J.R., P.K. Bhartia, O. Torres, C. Hsu, C. Seftor and E. Celarier (1997), Global Distribution of UV-absorbing aerosols from Nimbus7/TOMS data, J. Geophys. Res., 102, 16911–16922.

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Hsu, N.C., J.R. Herman, P.K. Bhartia, C.J. Seftor, O. Torres, A.M. Thompson, J.F. Gleason, T.F. Eck, and B.N. Holben (1996), Detection of biomass burning smoke from TOMS measurements, Geophys. Res. Lett., 23, 745–748. Klenk, K.F., P.K. Bhartia, A.J. Fleig, V.G. Kaveeshwar, R.D. McPeters, and P.M. Smith (1982), Total ozone determination from the Backscattered Ultraviolet (BUV) experiment, J. Appl. Met., 21, 1672–1684. McPeters, R.D., et al. (1996), Nimbus7 Total Ozone Mapping Spectrometer (TOMS) Data Product’s User’s Guide, NASA Reference Publication 1384, National Aeronautics and Space Administration, Washington, DC. Rodgers, C.D. (2000), Inverse methods for Atmospheric Sounding: Theory and Practice, World Scientific Publishing Co. Ltd. Torres, O. and P.K. Bhartia (1999), Impact of tropospheric aerosol absorption on ozone retrieval from backscattered ultraviolet measurements, J. Geophys. Res., 104, 21569–21577. Valks, P.J.M., J.F., de Haan, J.P. Veefkind, and R.F., van Oss, TOGOMI (2004), An improved total ozone retrieval algorithm for GOME, Proceeding Quadrennial Ozone Symposium, 1–8 June 2004, Kos, Greece. Wellemeyer, C.G., S.L. Taylor, C.J. Seftor, R.D. McPeters and P.K. Bhartia (1997), A correction for total ozone mapping spectrometer profile shape errors at high latitude, J. Geophys. Res., 102, 9029–9038. Wellemeyer, C.G., P.K. Bhartia, S.L.Taylor, W. Qin, and C. Ahn (2004), Version 8 total ozone mapping spectrometer (TOMS) algorithm, Proceeding Quadrennial Ozone Symposium, 1–8 June 2004, Kos, Greece. Yang, K., P.K. Bhartia, C.G. Wellemeyer, W. Qin, R.J.D. Spurr, J.P. Veefkind, J.F. De Haan (2004), Application of spectral fitting method to GOME and comparison with OMI DOAS and TOMS-V8 total ozone, Proceeding Quadrennial Ozone Symposium, 1–8 June 2004, Kos, Greece.

Chapter 4 The EOS Aura Mission* MARK R. SCHOEBERL

The Earth Observing System Missions and Aura The Earth Observing System (EOS) Aura satellite mission launched July 15, 2004. The Aura mission will make significantly improved measurements atmospheric constituent. Aura is designed to attack three science questions: (1) Is the ozone layer recovering as expected? (2) What are the sources and processes that control tropospheric pollutants? (3) What is the quantitative impact of constituents on climate change? Aura will answer these questions by globally measuring a comprehensive set of trace gases and aerosols (Table 1) at high vertical and horizontal resolution. Figure 1 shows the Aura spacecraft and its four instruments. The EOS Program consists of three core satellites, Terra (http://eosam.gsfc.nasa.gov/), Aqua (http://eos-pm.gsfc.nasa.gov/), and Aura (http://eosaura.gsfc.nasa.gov/) as well as several smaller satellites. Aura (Latin for “breeze”) was launched into an ascending node 705 km sun-synchronous polar orbit with a 98 ° inclination with an equator-crossing time of 13:45±15 minutes. The design life is five years with an operational goal of six years. Aura will fly in the same orbit track about 15 minutes behind Aqua, the Cloud-Aerosol Lidar and Infrared Pathfinder Satellite Observation (CALIPSO, http://www. calipso.larc.nasa.gov/) and Cloudsat launched on 28 April 2006 [Stephens et al., 2002, http://cloudsat.atmos.colostate.edu/]. This group of satellites, including the CNES PARASOL satellite (http://smsc.cnes.fr/PARASOL/GP_mission.htm) and the ESSP Orbiting Carbon Observatory (OCO, http://oco.jpl.nasa. gov/), is referred to as the “A-Train.” The measurements from Aura will be within 30 minutes of these other platforms. The near simultaneous measurements by the A-Train satellites will improve estimates of aerosol and cloud interaction and their role in precipitation and climate forcing.

*

This is a slightly modified version a paper that appeared in EOS [Schoeberl et al., 2004].

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FIGURE 1. Computer model of the Aura spacecraft showing the location of HIRDLS, MLS, OMI and TES. See Table 1. (Graphic by Jesse Allen, NASA Earth Observatory)

Science Objectives of the Aura Mission When combined with field campaign data, other satellite measurements, and ground-based observations, Aura measurements will provide unprecedented insights into atmospheric chemical and dynamical processes.

Is the Ozone Layer Recovering as Expected? Total Ozone Mapping Spectrometer (TOMS). Observations from 1978 show strong secular decrease in column ozone at extra-tropical latitudes. Although the Antarctic ozone hole area growth has slowed, significant late winter ozone depletions have now occurred in the Arctic [WMO, 2002]. Upper Atmosphere Research Satellite (UARS) data show a flattening in the stratospheric chlorine concentrations [Anderson et al., 2000]. A decrease in chlorine should lead to recovery of the ozone layer, but this recovery may be altered because of increasing greenhouse gas cooling [e.g., Shindel and Grewe, 2002]. As a result of the uncertainty in stratospheric trace gas trends, temperatures and dynamical feedback processes, current models used to assess the ozone layer do not agree on the timing of the ozone layer recovery [WMO, 2002]. The stratospheric measurements made by Aura will permit a complete assessment of the chemical processes controlling ozone. First, high vertical resolution ozone profiling by MLS and HIRDLS will provide the best information ever on ozone change. Second, five of the major radicals that participate in ozone destruction (ClO, OH, HO2, BrO, and NO2) will be measured by either HIRDLS or MLS. Third, MLS and HIRDLS also measure the important reservoir gases, HCl, ClONO2, and HNO3. Fourth, the Aura instrument payload will make measurements of the long-lived source gases including N2O, H2O, CH4, CFC’s. Finally,

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TABLE 1. Aura Instruments, Principle Investigators and Measurements (Most measurements will have a precision of 10% or better.) Acronym

Name

Instrument PI

Constituent

HIRDLS

High Resolution Dynamics Limb Sounder

Profiles of T, O3, H2O, CH4, N2O, NO2, HNO3,, N2O5, CF3Cl, CF2 Cl2, ClONO2, Aerosols

MLS

Microwave Limb Sounder

John Gille, National Center for Atmospheric Research & U. of Colorado; John Barnett, Oxford University Joe Waters, Jet Propulsion Laboratory

OMI

Ozone Monitoring Instrument

Pieternel Levelt, KNMI, Netherlands

TES

Tropospheric Emission Spectrometer

Reinhard Beer, Michael Gunson, Jet Propulsion Laboratory

Profiles of T, H2O, O3, ClO, BrO, HCl, OH, HO2, HNO3, HCN, N2O, CO, Cloud ice. Column O3, SO2, aerosols, NO2, BrO, OClO. HCHO, UV-B, cloud top pressure, O3 profiles. Profiles of T, O3, NO2, CO, HNO3, CH4, H2O.

Instrument Description Limb IR filter radiometer from 6.2µ to 17.76µ 1.2 km vertical resolution up to 80 km.

Microwave limb sounder 118 GHz to 2.5 THz 1.5–3 km vertical resolution Hyperspectral nadir imager, 114 ° FOV, 270-500 nm, 13 × 24 km footprint for ozone and aerosols Limb (to 34 km) and nadir IR Fourier transform spectrometer 3.2–15.4µ Nadir footprint 5.3 × 8.5 km, limb 2.3 km

OMI will continue the TOMS/SBUV global column and profile ozone trend measurements.

What Are the Sources and Processes That Control Tropospheric Pollutants? Tropospheric ozone production occurs when CO, volatile organic compounds (VOCs), and nitrogen oxides are exposed to sunlight. These ozone precursors are directly linked to urban sources, and the atmosphere can transport both ozone and its precursors over large distances. The Aura mission is designed to produce the first global assessment of tropospheric ozone and its precursors, as well as assess the stratospheric contribution to the tropospheric ozone budget. The measurements from TES as well as column measurements from OMI combined with

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stratospheric measurements from HIRDLS and from MLS will provide new information on the pollution sources and transport.

What Is the Quantitative Impact of Aerosols and Upper Tropospheric Water Vapor and Ozone in Climate Change? One of the sources of uncertainty in climate change comes from our poor understanding of the changing concentration of upper troposphere and lower stratospheric (UT/LS) constituents [Houghton et al., 1995]. On decadal time scales, climate signals from greenhouse gas changes and changes in reactive constituents are intertwined. Aura’s MLS and HIRDLS instruments have been specifically designed to study the UT/LS in unprecedented detail. MLS can make constituent measurements through thin clouds while HIRDLS can scan vertically across the limb over a wide horizontal range through cloud gaps. The issues of climate radiative forcing cannot be solved by Aura alone. Data from Aqua, Cloudsat, and CALIPSO will also be needed to address climate issues fully.

Aura Instrument Descriptions Figure 2 shows the instrument fields of view. MLS will make limb sounds observing forward. OMI and TES will make nadir soundings. HIRDLS and TES will make limb soundings observing backward. The advantage of this instrument configuration is that each of the instruments can observe the same air mass within ~13 minutes.

FIGURE 2. Aura instrument instantaneous fields of view, looking toward the back of the spacecraft. MLS limb measurements, green, OMI nadir measurements, blue swath. TES limb and nadir measurements, red; HIRDLS measurements yellow. Limb sounders scan vertically. (Graphic by Jesse Allen, NASA Earth Observatory.)

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HIRDLS HIRDLS is an infrared (6–1 µm) limb scanning, 26-channel filter radiometer. HIRDLS makes temperature and trace gases observations from the UT/LS to the mesosphere (Table 1) [Gille et al., 2003]. HIRDLS was jointly built by the United Kingdom and the United States. HIRDLS can also detect clouds and thus determine the altitude of polar stratospheric clouds and tropospheric cloud tops. HIRDLS has higher horizontal resolution than any previous limb sounder through the use of a programmable azimuth scan in conjunction with a rapid elevation scan. Special observing modes can be used to observe geophysical events like volcanic clouds.

MLS MLS is a 118 GHz–2.5 THz, limb scanning microwave emission spectrometer that measures temperature and constituents from the UT/LS to the mesosphere (Table 1) [Waters et al., 1999]. MLS also has capability of measuring upper tropospheric water vapor in the presence of cirrus. Aura MLS is based on UARS MLS but uses more advanced technology to make additional constituent measurements. These new constituents play an important role in stratospheric chemistry (i.e., HCl, N2O, OH, HO2 and BrO). The UARS MLS was able to measure upper tropospheric water vapor [Read et al., 2001], which is essential for understanding climate variability. Aura MLS will improve this measurement.

OMI The OMI instrument is a contribution of the Netherlands’ Agency for Aerospace Programs (NIVR) in collaboration with the Finnish Meteorological Institute (FMI) [Levelt et al., 2000]. OMI will continue the TOMS record for total ozone and other atmospheric parameters related to ozone chemistry and aerosols. OMI is a 0.24–0.50-µm, visible-UV 740 band, cross track hyperspectral imager. OMI will provide global coverage in one day at 13 × 24 km spatial resolution. A combination of backscatter ultraviolet retrieval algorithms and forward modeling will be used to generate the various OMI data products [Ahmad et al., 2003].

TES TES is a 3.2–15.4 µm, high-resolution infrared-imaging Fourier Transform spectrometer. TES has a spectral resolution of 0.025 cm−1 measuring most of the radiatively active molecular species in the Earth’s atmosphere [Beer et al., 2001]. TES can make both limb and nadir observations. TES can target within 45 ° of the local vertical, or produce regional transects over clear regions up to 885 km in length. TES will provide global maps of trace gases listed in Table 1. Because TES measures the entire IR spectrum, the potential exists to retrieve a large number of other gases (e.g., ammonia); although the retrieval of these gases will be done in a research mode.

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Aura Instrument Synergy Because Aura instruments observe the same air mass within ~13 minutes, it will also be possible to better interpret photochemical processes involving constituents measured simultaneously as has been done with UARS data (e.g., Douglass et al., 1995). In addition, HIRDLS and MLS limb sounding will provide ozone profiles that can be combined with the OMI column ozone observations. It will then be possible to separate the stratospheric component of the column ozone and thus estimate the tropospheric ozone column. The tropospheric column can be compared with the TES direct measurement of tropospheric ozone.

Summary The five-year EOS Aura mission will provide a significantly improved tropospheric and stratospheric constituent measurements. Although there are only four instruments on Aura, the breadth of the their capability combined with the other A-Train measurements will provide a powerful tools to attack future questions about changing atmospheric composition and its impact on climate. The instruments OMI, MLS, and TES are on and performing nominally. Observations of the 2004 Antarctic ozone hole have already been made by MLS and OMI. HIRDLS is experiencing radiative anomalies due to a piece of kapton blocking the aperture; its future performance is uncertain.

References Anderson, J.J., et al., 2000: Halogen occulation experiment confirmation of stratospheric chlorine decreases in accordance with the Montreal protocol, J. Geophys. Res., 105, 4483–4490. Ahmad, S.P., et al. Atmospheric products from the ozone monitoring instrument (OMI), Proceedings of SPIE July 2003, San Diego, U.S.A., 5151–66, 2003. Beer, R., T.A. Glavich and D.M. Rider, Tropospheric emission spectrometer for the Earth Observing System’s Aura satellite. Applied Optics, 40, 2356–2367, 2001. Douglass, A.R. et al., Interhemispheric Differences in Springtime Production of Vortex Chlorine Reservoirs, J. Geophys. Res., 100, 13967–13978, 1995. Gille, J., et al., A. Lambert and W. Mankin, The High Resolution Dynamics Limb Sounder (HIRDLS) Experiment on Aura, Proc. of SPIE, 5152, 162–171, 2003. Houghton, J.T., et al. (eds.), Climate Change 1994, The IPCC Scientific Assessment, Cambridge University Press, 339 pp., 1995. Levelt, P.F., et al. Science Objectives of EOS-AURA’s Ozone Monitoring Instrument (OMI), Proceedings of the Quadrennial Ozone Symposium, Sapporo July 2000, pp. 127–128, 2000. Read, W.G., et al., UARS Microwave Limb Sounder Upper Tropospheric Humidity Measurements: Method and Validation. J. Geophys. Res., 106, 32,207–32,258, 2001. Schoeberl et al. Earth Observing Systems Benefit Atmospheric Research, EOS, 85, 177–178, 2004.

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Shindell, D.T. and V. Grewe, Separating the influence of halogen and climate change on ozone recovery in the upper stratosphere, J. Geophys. Res., 107, doi:10.1029/2001 JD000420, 2002. Stephens, G.L., et al., The CloudSat mission and the A-train—A new dimension of spacebased observations of clouds and precipitation, Bull. Amer. Met. Soc., 83, 1771–1790, 2002. Waters, J.W., et al., The UARS and EOS Microwave Limb Sounder (MLS) Experiments. J. Atmos. Sci., 56, 194–218, 1999. World Meteorological Organization, Scientific Assessment of Ozone Depletion: 2002, Report No. 47, 2002.

Chapter 5 MIPAS Experiment Aboard ENVISAT HERBERT FISCHER

Introduction At March 1, 2002, the European Space Agency (ESA) launched the ENVISAT satellite in a polar, sun-synchronous orbit. The large satellite with a mass of more than eight tons is orbiting in a height of about 800 km with a descending node of 10:00 h. From the ten experiments on board three are used to explore the composition of the atmosphere. One of them is the Michelson Interferometer for Passive Atmospheric Sounding (MIPAS). The Institute for Meteorology and Climate Research (IMK) has developed in the past several MIPAS instruments that have been used in ground stations, aboard aircrafts and balloon gondolas [Fischer, 1993]. Finally, the IMK has proposed a corresponding satellite experiment to ESA who has included this MIPAS experiment in the core payload of ENVISAT (see Figure 1).

MIPAS/ENVISAT Experiment MIPAS is a high-resolution Fourier Transform Spectrometer that measures the atmospheric limb emission simultaneously in four spectral bands in the middle infrared [H. Fischer et al., 2000]. The whole spectral interval covers the range from 4.15 µm to 14.6 µm. The unapodized spectral resolution is 0.035 cm−1, which allows the detection of many single spectral emission lines. MIPAS is continuously performing measurements during day and night, providing each day a full coverage of the globe. During an orbit the instrument performs 75 limb scans and in addition calibration measurements when operating in the nominal mode. The instrument is very flexible, i.e., it can also be used for performing a considerable number of other observation modes with various scientific objectives. By analysing a limb scan sequence vertical profiles of temperature and many trace gases can be derived. Furthermore, these data are appropriate to determine information about aerosols as well as particles of Polar Stratospheric and cirrus clouds.

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FIGURE 1. Optical design of the MIPAS/ENVISAT experiment.

Combining the profiles along an orbit, the distribution of the atmospheric parameters can be presented as a function of altitude and latitude. The MIPAS instrument is a dual-slide, dual-port interferometer that is using all the incoming radiance for measurement and is in various instrumental parts redundant. Figure 1 shows the optical design of MIPAS-ENVISAT including the azimuth and elevation scan unit, the telescope, the interferometer, and the focal plane subsystem with the detectors. Table 1 summarizes the main specifications of the MIPAS experiment. TABLE 1. Summary of MIPAS Specifications ENVISAT launch: 1st March 2002 Michelson interferometer Spectral range: 4.1–14.6 µm Spectral resolution: 0.035 cm−1 Altitude range: Nominal mode: 6–68 km several other observation modes Global coverage: measurements from pole to pole Measurement schedule: continuous Measurement parameters: -p, T, H2O, O3, CH4, N2O, HNO3, NO2 -NO, N2O5, HNO4, CIONO2, CIO, CO, CFCs, NH3, C2H6 HDO, O3 isotopomers and others (more than 25 parameters)

operational processing scientific processing

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The enormous amount of information in one MIPAS spectrum is shown in Figure 2. This emission spectrum of the atmosphere associated to a tangent altitude of 26 km (definition see below in Fig. 4) impresses by the large number of spectral lines which stem from many atmospheric trace gases as indicated in the figure. Most prominent are the absorption bands of CO2 (15 µm), O3 (9.6 µm), CH4 (7.7 µm), and H2O (6.3 µm).

MIPAS-Experiment Aboard ENVISAT Data Processing An overview of the measurement and data processing is given in Figure 3. The interferometer transforms the atmospheric radiance in an interferogram, which has to be converted back to a spectrum by applying a Fourier transformation. In the following this spectrum has to be calibrated by using blackbody and deep space measurements. A further step in data processing is the inversion of spectral radiances in trace gas concentrations. A trace gas profile can be derived from a sequence of limb measurements. The limb sounding geometry is shown in Figure 4. The ray path is mainly defined by the tangent altitude, which corresponds to the minimum distance between ray path and Earth surface. The information about the vertical structure of the atmosphere is attained by scanning across the atmosphere. The Radiative

FIGURE 2. MIPAS spectrum for a tangent height of 26 km showing the enormous amount of information contained.

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FIGURE 3. Overview of the measurement and data processing.

FIGURE 4. Limb sounding geometry and the Radiative Transfer Equation.

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Transfer Equation (RTE) describes the connection between the measured radiances and the distribution of the atmospheric parameters. A corresponding computer code is called KOPRA (Stiller et al., 2000), which has been developed at the IMK. It is based on the up-to-date knowledge about the radiative processes in the atmosphere including the ray tracing for an elliptical earth, the consideration of horizontal inhomogeneities, and non-LTE (local thermodynamic equilibrium) effects (Table 2). In order to determine vertical profiles of atmospheric parameters from the radiance measurements, the RTE has to be inverted (i.e., the integral equation has to be solved). Methods for this so-called retrieval are described by Rodgers [2000]. We are using the optimal estimation method with a Tikhonov smoothing operator. A sequence of limb emission measurements (between 68.4 km and 6.4 km tangent altitude) is shown in Figure 5. First of all, a gap of information can be recognized in the spectral range between 970 and 1020 cm−1. At high altitudes the emitted radiance is relatively weak; only the CO2-band below 760 cm−1 and the O3-band around 1050 cm−1 are clearly visible. With decreasing tangent altitude the emitted radiance becomes stronger and additional absorption bands show up, e.g., the CO2 laser band around 960 cm−1, the HNO3-bands around 875 cm−1, and the CFC bands around 850 cm−1 and 920 cm−1. The HNO3 bands grow rapidly in the lower stratosphere due to the shape of the vertical profile of the HNO3 mixing ratio. In the upper troposphere the spectrum exhibits some single strong lines in the range of the atmospheric window between 800 and 1000 cm−1. They are caused by weak water vapor lines becoming visible in the upper troposphere due to the quickly increasing water vapor mixing ratio.

TABLE 2. Main Characteristics of the Radiative Transfer Code KOPRA Main Characteristics of KOPRA ● Currently ~70 different trace species, many of them relevant in the UTLS only ● Non-LTE radiative transfer modelling ● Vertical variation of isotopomeric abundances ● Accurate modelling of spectroscopic features (CO , line mixing, self/air broadening, pressureshift, 2 line shapes, continua, p/T-dependent cross-sections of heavy molecules, ...) ● Particle-caused absorption, emission and single scattering ● Pseudo-analytic derivatives ● Generic non-LTE state population model coupled into KOPRA ● Mie model for calculation of absorption and scattering coefficients and phase function coupled into KOPRA ● Raytracing for elliptical earth ● Consideration of horizontal inhomogeneities (P, T, vmr, ...) by 3-D fields ● Derivatives wrt —Chemical composition and microphysical properties of particles (size distribution, ...) —Non-LTE process parameters (collision rates, production rates, ...) are provided ● Modelling of instrument characteristics (FOV, ILS, ...) ● User-defined numerical accuracy

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FIGURE 5. Sequence of limb emission spectra in the altitude region between 68.4 km and 6.4 km.

Derivation of Trace Gas Distributions As mentioned above, from the sequences of measured limb radiance spectra vertical profiles of trace gases can be derived. Combining the vertical profiles along the satellite orbit allows to generate cross sections of atmospheric parameters. In the following examples of results will be shown which prove the enormous capability of the MIPAS experiment. Figure 6 shows stratospheric cross sections of long-lived trace gases, potential vorticity (PV) and O3 along the satellite orbit from 40 °S across the South Pole to 40 °S. The three columns are related to different dates: 20 September, 24 September, and 13 October 2002 (from left to right). The time period is connected with the break-up of the Polar Vortex over Antarctica. At 20 September 2004 the Polar Vortex is seen very clearly by the low mixing ratios of N2O and CH4 in the lower stratosphere. The potential vorticity (PV) as derived from ECMWF-data confirms this statement. Simultaneously, low ozone values are obviously existing over Antarctica. The split up of the polar vortex is recognized at 24 September 2002 connected with high ozone concentrations at high southern latitudes. At 13 October 2002 a re-establishment of the polar vortex has happended. On the other hand, this polar vortex is much less pronounced in the middle stratosphere as can be seen in the CH4 cross section. This example shows clearly the

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FIGURE 6. Stratospheric cross sections of trace gases and potential vorticity along the satellite orbit from 40 °S across the South Pole to 40 °S for three different dates.

comprehensive information content of the MIPAS data about the stratospheric dynamics [see also Glatthor et al., 2004]. The ozone depletion in the southern polar vortex in spring time can be demonstrated by intercomparing the ozone with the long-lived CFC-11 mixing ratio inside and outside the polar vortex (see Figure 7). At 20 September the O3 mixing ratio outside the polar vortex for CFC-11 values of 0.02 to 0.01 ppbv is much higher than the corresponding O3 mixing ratios inside the polar vortex. In case of mid-October the ratio of O3 mixing ratio outside to inside is even greater than before, i.e., the ozone depletion was ongoing during end of September and beginning of October 2002. Essential for the stratospheric ozone depletion is the concentration of ClO in the atmosphere; under high ClONO2 concentrations this process is not effective. MIPAS is capable to measure also ClO under disturbed chemical conditions as in the polar winter/spring. Figure 8 shows the stratospheric cross sections of ClO and ClONO2. High amounts of ClO are only visible in the polar lower stratosphere at the day side. ClONO2 shows the typical collar structure over Antarctica [see Höpfner et al., 2004a]. Results on validation of the trace gas distributions have been described in the cited articles. For the main trace gases there is an overview article available based on the ENVISAT validation workshop in May 2004 in Frascati (see http://envisat.esa.int, Fischer H. and H. Oelhaf, ESA-proceedings).

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O3 versus CFC-11 10 outside

inside

O3 (ppmv)

0 0

CFC-11 (ppbv)

0.4

20-21 Sep (red) and 11-13 Oct (blue)

FIGURE 7. Intercomparison of ozone mixing ratios inside and outside the polar vortex for 20/21 September and 11/12/13 October 2002.

MIPAS CIONO2 and CIO: inorganic chlorine on September 20, 2002, orbit 2910

CIONO2: Envisat orbit 2910, 20020920 09:47-10:36 UT

[ppbv]

CIO: Envisat orbit 2910, 20020920 09:47-10:36 UT

30

30

25

25

2.5

Altitude [km]

2.0

20 15 10 10

1.5 1.0

20

0.5

15

30

700 K 625 K 550 K 475 K 400 K

60 90 60 Southern latitude [deg]

30

10

10 10

0.0 −0.5

30

60 90 60 Southern latitude [deg]

30

10

collar structure

FIGURE 8. Stratospheric cross sections of ClO and ClONO2 for higher latitudes in the southern hemisphere at 20 September 2002.

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Determination of Cloud Parameters MIPAS is a versatile instrument; from the measurements not only trace gas concentrations but also parameters of thin clouds can be derived. Höpfner et al. [2004b] have proven that high resolution broad-band IR spectra can be used to determine the mean size and the volume density of the cloud particles as well as their chemical composition. The limb spectra of Polar Stratospheric Clouds (PSC) show very special signatures, which are caused by a superimposition of cloud emitted radiance and scattered tropospheric radiance (Figure 9). The interpretation of the spectra as measured along different orbits across the Antarctic region has yielded spatial distributions of PSCs for different dates (see Figure 10). These results show clearly that the development of the PSCs is connected with complex processes in the atmosphere depending on temperature, dynamics, and chemical composition. Höpfner (personal communication) has already proven that thin cirrus clouds in the tropopause region exhibit the same typical spectral signatures as PSCs, i.e., also size distribution and volume density of these ice particles can be derived.

FIGURE 9. Superimposition of PSC emitted radiance and scattered tropospheric radiance for a limb observation.

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FIGURE 10 Spatial distribution of different types of PSCs over Antarctica at four dates.

Conclusion and Outlook MIPAS provides high-quality global measurements during day and night of temperature, more than 25 trace constituents and parameters of thin clouds. The validation of the results yields generally good agreement with the measurements of various other experiments. The major warming over the Antarctic stratosphere in September/October 2002 was well covered with MIPAS measurements. Firsttime ClO concentrations have been detected with IR instruments from space. Simultaneous observations of trace gas concentrations and cloud parameters (PSCs, cirrus clouds) were performed. Future MIPAS measurements have to be carried out with lower spectral resolution (about 40%) because the interferometer has a problem with the retroreflector slides. On the other hand, the lower spectral resolution does not reduce the information content of the measurements significantly. Acknowledgement The versatile contributions of several ESA colleagues, members of the MIPAS Scientific Advisory Group, and many scientists of my institute are acknowledged.

References Fischer, H., (1993) Remote sensing of atmospheric trace gases. Interdisciplinary Science Reviews 18: 185–191.

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Fischer, H., C. Blom, H. Oelhaf, B. Carli, M. Carlotti, L. Delbouille, D. Ehhalt, J.-M. Flaud, I. Isaksen, M. Lopez-Puertas, C.T. McElroy, and R. Zander. (2000): EnvisatMIPAS, the Michelson Interferometer for Passive Atmospheric Sounding; An instrument for atmospheric chemistry and climate research, ESA SP-1229, C. Readings and R.A. Harris, eds. (European Space Agency, Noordwijk, The Netherlands, 2000). Glatthor, N., et. al. (2004), Mixing processes during the Antarctic vortex split in September/October 2002 as inferred from source gas and ozone distributions from MIPAS/ENVISAT, J. Atmos. Sci., Special issue on Antarctic Vortex 2002, accepted January 2004. Höpfner, M., et. al. (2004a), First spaceborne observations of antarctic stratospheric ClONO2 recovery: austral spring 2002, J. Geophys. Res., 109, D11308. Höpfner, M. (2004b), Study on the impact of polar stratospheric clouds on high resolution mid-IR limb emission spectra, J. Q. S. R. T., 83, 93–107. Rodgers, C.D., (2000), Inverse Methods for Atmospheric Sounding: Theory and Practice, Vol. 2, edited by F.W. Taylor, World Scientific, River Edge, N.J. Stiller, G.P., (ed.), (2000), The Karlsruhe optimized and precise radiative transfer algorithm (KOPRA), Wiss. Ber. FZKA 6487, Forschungszentrum Karlsruhe, Germany.

Part II Aircraft and Ground-Based Intensive Campaigns

Chapter 6 Probing the Atmosphere with Research Aircraft—European Aircraft Campaigns ULRICH SCHUMANN

Why Aircraft Campaigns? Aircraft are used to investigate the atmosphere and the ground by visual observations and by measurements with instruments onboard the aircraft. They allow for ●

●

● ● ●

●

●

Targeted measurement of atmospheric properties above ground with in-situ instruments and with airborne remote sensing instruments along a given flight path (e.g., from 100 m up to 20 km; looking up and down) Measurements at high temporal/spatial resolution (order 1 s, 10 m, as often required) Data for process studies (e.g., cloud formation) Data for validation of models (e.g., chemical transport models) Test and validation of remote sensing (e.g., GOME, SCIAMACHY, MIPAS, e.g., Heland et al., 2002) Measurements of otherwise not measurable properties (e.g., tropospheric NO, turbulence, cloud condensation nuclei) Simultaneous measurements of many parameters (meteorology, air composition, aerosols, cloud physics etc.)

This lecture is based on experiences gained within several European research projects with research aircraft, such as POLINAT, EULINOX, INCA, and TROCCINOX. The projects dealt with — POLINAT: Pollution in the North Atlantic tropopause region induced by aircraft and other sources [Schlager et al., 1997, 1999; Schumann et al., 1995, 2000; Ziereis et al., 1999, 2000], — EULINOX: Lighting induced nitrogen oxides over Europe [Höller et al., 1999; Huntrieser et al., 2002], — INCA: Impact of different aerosol and trace gas concentrations and atmospheric dynamics on cirrus clouds in the Southern and Northern hemisphere [Ström et al., 2003; Baehr et al., 2003; Minikin et al., 2003; Ovarlez et al., 2002; Ziereis et al., 2004; Gayet et al., 2002, 2004], and 85

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— TROCCINOX: Nitrogen oxides from lightning (besides cirrus and aerosol properties) near deep convective clouds in a tropical continent [Schumann et al., 2004]. Instead of describing the scientific results of these and related projects, this short contribution concentrates on the European research aircraft fleet and the methodological aspect of airborne measurements.

European Research Aircraft The European fleet for airborne research is, to a large extent, organized within the EU-network EUFAR; see www.eufar.net. Within this network, the European Commission of the European Union provides financial support for transnational access (preferably for young scientists from countries outside the country of the aircraft operator). The EUFAR fleet composes presently 22 aircraft within 5 categories. Figure 1 lists the aircraft by type and operator together with a small photo. In addition to these research aircraft, in Europe several instrumented airliners are used, in particular several Airbus A340 aircraft within the project MOZAIC (with measurements of H2O, O3, CO, NOy; see paper by J.-P. Cammas in this volume), and a Lufthansa A340-600 (previously LTU B767) carrying a container with a large suite of instruments within the project CARIBIC [Brenninkmeijer et al., 1999]. In 1995–1997, a Swiss Air B747 was equipped with a container carrying instruments to measure NO, NO2, O3 within the project NOXAR [Brunner et al., 2001]. The Swiss project NOXAR was coordinated with the EU-Project POLINAT 2 [Schumann et al., 2000]. In the near future, it is expected to get access to a new research aircraft known under the project acronym HALO. Similar to the project HIAPER in the USA, HALO will be an instrumented Gulfstream G550. The HALO investment is funded by the German Ministry of Research and Education (BMBF) and by German research institutions (MPG, HGF). HALO will be operated by DLR under an open consortium including in addition the German Science Foundation (DFG), the Leibniz Society (WGL), and other institutes. HALO is expected to become operational in 2008. Compared to other existing European research aircraft, HALO will go “higher, farther, larger:” up to 15.5 km altitude, more than 8000-km range, and up to 3000-kg payload.

Instrumentation Some of the research aircraft are equipped with a specific set of instruments, which can only with large efforts be changed into new configuration. However, many of the research aircraft are equipped with devices (racks, openings, inlets, windows, power supply, data recording system, etc.) that allow for equipment

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Stratospheric Jet Geophysica (Geophysica EEIG)

High Level Jets Falcon 20 (DLR) Citation II (NLR) Falcon 20 (SAFIRE) Learjet (Enviscope)

Large Aircraft ATR-42 (SAFIRE) BAe 146 (Met Office)

FIGURE 1. The EUFAR fleet.

with variable sets of instruments. For example, the Falcon research aircraft of DLR was equipped with instruments as listed in Table 1, during the field campaign of the project TROCCINOX in Jan–March 2004. In particular, DLR has developed instruments to be flow on the Falcon for trace gas measurement systems [Schlager et al., 1997], particle instrumentation [Schumann et al., 2002; Minikin et al., 2003], Lidar instruments [Ehret et al., 1999, 2000; Fix et al., 2002] (available for O3, H2O, wind profiles), and standard instruments (position, speed, temperature, humidity, wind etc.) [Schumann et al., 1995]. The trace gas measurement system is available to measure many of the components important for photochemical reactions in the troposphere.

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Medium-Sized Aircraft Do 228-101 (NERC) Cessna Caravan (DLR) CASA 212 (INTA, 2 aircraft) Do 128 (TU Braunschweig) Twin Otter (GSF) Cessna Caravan (GSF) King-Air 200 (TAU)

Low-Level Aircraft Piper Aztec (SAFIRE) Partenavia (Enviscope) Cessna 207 (FU Berlin) Dimona (MetAir) ASK 16 (TU Berlin) ENDURO (IFU) Skyarrow (ISAFM)

FIGURE 1. (Continued)

In addition to the particle instrumentation listed in Table 1, an extended set of instruments was available within the project INCA, to cover the whole size range from a few nanometers to millimeters; see Figure 2. During TROCCINOX, the Lidar was flown looking upwards to measure Lidar backscattering—at wavelengths 532 and 1064 nm—and H2O by differential absorption with two wavelengths near 935 nm, in the altitude range up to 16 km; see Figure 3 for an example. During transit from Germany to Brazil and back, the Lidar was looking downward to measure aerosol and H2O profiles from 11 km downward to the ground within the troposphere. In other projects, the Lidar systems were flown to measure O3 and wind profiles.

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TABLE 1. Falcon Instrumentation for TROCCINOX Technique DIAL DIAL upward or downward UV absorption Chemiluminescence Chemiluminescence + Au converter VUV fluorescence Condensation particle size analyzer (lower cutoff 4, 9, 13 nm) Nonvolatile CN CN counter with heated inlet (lower cutoff 13 and 50 nm) Aerosol absorption PSAP (particle soot absorption coefficient photometer) Aerosol and cloud PCASP-100X (0.2–1.0 µm) element size distribution FSSP-300/300 (0.3–20/5–100 µm) Position, wind INS, GPS, 5-hole probe

∆t 5 sec 1 min 5 sec 1 sec 1 sec 5 sec 1 sec

Precision/Accuracy 1/5% 5/10% 1/5% 3%/10% 5%/15% 1/3 nmol/mol 5/10%

1 sec

5/10%

30 sec (BL) 20–60 sec 1 sec

10/20%

Temperature Humidity NO2 photolysis, J(NO2)

1 sec 1 sec 1 sec

Species/Parameter Aerosol Profile Remote H2O Profile Ozone NO NOy CO Condensation nuclei (CN)

Rosemount Lyman-alpha Filter radiometer

10/20% 200 m (horiz.) 0.05/1 m/s (h) 0.05/0.3 m/s (v) 0.1/0.5 K 0.01/0.3 g/m3 1 E-4/5 E-4

Number concentration, cm−3

The spatial and temporal resolution of airborne measurements is usually far higher than what can be achieved by remote sensing -on the other hand, the coverage reached by satellite instruments is far higher than for aircraft. For properties influenced by turbulence or clouds in the atmosphere and for short lived chemical species, high resolution is particular important. Typically the instruments reach a temporal resolution of about 1 s. Other important properties of instruments are precision and accuracy, i.e., the relative and absolute resolution of the measured parameter. See Table 1 for the properties of the Falcon instrumentation during TROCCINOX. 105

Aerosol Particles

103

Condensation Particle Counter Aerosol Volatility Analyzer

101 10−1 10−3 10−5

Cloud Elements

PCASP 100X FSSP 300 FSSP 100 ER Polar Nephelometer 2 DC

0.01

0.1 1 10 100 Particle diameter, µm

1000

FIGURE 2. Typical aerosol size spectrum and its coverage by various instruments. [A. Petzold, A, Minikin, DLR, pers. Communication, 2003.]

Ulrich Schumann

1.00 1.72 2.96 5.10 8.77 15.1 26.0 44.7 77.0 132 227 392 675 1162 2000

10.0 13.9 19.3 26.8 37.3 51.8 72.0 100 138 193 268 372 517 719 1000

Backscatter Ratio at 532 nm (II Polarisation)

Altitude [km]

18 16 14 12 10 8

18:45:00

18:50:00

18:55:00 Time (UTC)

19:00:00

19:05:00

Backscatter Ratio at 1064 nm (II Polarisation)

18 Altitude [km]

1.00 1.53 2.35 3.61 5.54 8.50 13.0 20.0 30.7 47.1 72.2 110 169 260 400

16 14 12 10 8

0

50

100

150 200 Distance [km] Water Vapour Mixing Ratio [ppm]

18 Altitude [km]

90

250

16 14 12 10 8

−17.5

−17.0

−16.5 Latitude [⬚N]

−16.0

−15.5

FIGURE 3. Observations of backscatter and water vapor near the anvil of a tropical thunderstorm during TROCCINOX [G. Ehret, A. Fix, H. Flentje, M. Wirth, personal communication, 2004].

Quality Checks An important task in performing airborne measurements are quality checks of the instrumentation. Such checks include many activities such as model analysis and laboratory studies, preflight, inflight, and postflight tests and calibrations. One specific activity of large importance are intercomparison flights, where two aircraft fly close to each other, preferably wing by wing, to measure the same atmospheric parameter with their individual instruments. For example, during the summer 2000 Export aircraft campaign (European eXport of Precursors and Ozone by long-Range Transport), see Figure 4, two comprehensively instrumented research aircraft, the British C-130 and the German Falcon, measuring a variety of chemical species flew wing tip to wing tip for a period of one and a quarter hours [Brough et al., 2003]. During this interval a comparison was undertaken of the measurements of (NO), NOy, CO, and O3. The comparison was performed at two different flight levels, which provided a 10-fold variation in the concentrations of both NO (10 to 1000 pmol/mol) and NOy (200 to over 2500 pmol/mol). Large peaks of NO and NOy observed from the Falcon 20, which were at first thought to be from the exhaust of the C-130,

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FIGURE 4. Left. The MRF C-130 and the DLR Falcon 20 during EXPORT. Right: Flight paths of the Falcon (light gray) and the C-130 (dark gray) [Brough et al., 2003].

8000 2 1200 1000 800

DLR NOy DLR NO Altitude NO NOy

3

4

6000

1 4000

600 400

Altitude (Metres)

DLR and UEA NO, NOy Mixing Ratios, pptv

were also detected on the 4-channel NOx,y instrument aboard the C-130. See Figure 5, for example. These peaks were a good indication that both aircraft were in the same air mass and that the Falcon 20 was not in the exhaust plume of the C-130. Correlations and statistical analysis are presented between the instruments used on the two separate aircraft platforms. These were found to be in good agreement, giving a high degree of correlation for the ambient air studied. Any deviations from the correlations are accounted for in the estimated inaccuracies of the instruments. These results help to establish that the instruments aboard the separate aircraft are reliably able to measure the corresponding chemical species in the range of conditions sampled. An intercomparison of airborne in-situ water vapor measurements by two European research projects (MOZAIC and POLINAT) was performed from aboard the respective Airbus (MOZAIC) and Falcon (POLINAT) aircraft [Helten et al., 1999]. The intercomparison took place southwest of Ireland on September 24, 1997, at 239 hPa flight level. MOZAIC uses individually calibrated capacitive humidity sensors for the humidity measurement. POLINAT employs a cryogenic frost-point hygrometer developed for such measurements [Ovarlez et al., 2003].

2000 200 0 0 13:15 10/08/00

13:30

13:45

14:00

14:15

14:30

Time (Greenwich Mean Time, GMT)

FIGURE 5. The mixing ratios for NO and NOy observed at altitudes between 6000 and 8000 m, during the comparison over southern Germany [Brough et al., 2003].

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For conversion between humidity and mixing ratio, ambient temperature and pressure measurements onboard the respective aircraft are used. The Falcon followed the AIRBUS at a distance of 7 km to 35 km with a time-lag increasing from 30 s to 160 s. The water vapor volume mixing ratio measurements in the range of 80 to 120 ppmv of both instruments are in excellent agreement, differing by less than ±5%, where the trajectories of both aircraft are very close. However, the relative humidity (RH) calculated from POLINAT frost-point measurements and the Falcon -PT500 temperature sensor is up to 15% higher relative to the RH of MOZAIC. The agreement improved to within 5% when using the temperature measurement of the PT100 sensor instead of the temperature measurement of the PT500 sensor for RH determination of POLINAT.

Campaign Planning An important part of any airborne campaign is the campaign planning. This topic is usually not reported in the scientific literature. It simply belongs to the tasks of any such experiment. Nevertheless, it is important to recognize that this task usually takes months before the experiment, and its quality is one of the main reason for project success or failure. The “White Book” by Rob MacKenzie et al. (written as part of the TROCCINOX project), includes the following tasks: Pre-Campaign Phase: Instrument preparation and upgrading, model development, tools development for flight planning, and logistics and operational planning, aircraft preparation. Campaign Phase: Calibration of the instruments, campaign performance, data analysis Interpretation Phase: Modeling of the observed scenarios using state-of-the-art models and hypotheses tests by intercomparison with the data sets. The results of the Pre-Campaign Planning are documented in a document called “White Book”. Its content includes — Key scientific questions — Campaign methodology: Site selection, observational platforms, alternate airports, flight objectives, instrumentation — Flight templates — Timelines — Contingency — Management — Briefing Some of the ingredients of the White Book were presented in the workshop. The reader may refer to the project homepage for further details. An essential part of campaign organizations are recalculations of the chemical composition of the atmosphere, see Flatoy et al. [2000].

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Databases Usually, the results of the data of a measurement campaign are collected in a project specific database, which includes all the original measurements for access by all project partners. Usually, the data banks are pass-word-controlled until at least one year after the project ends. It is common practice to allow other users access on the condition that the results are published in agreement with the experimenters, including co-authorship were reasonable. The databases can usually be found via the homepages of the specific projects in the internet, e.g., http://www.pa.op.dlr.de/polinat/,../inca/,../eulinox/,./troccinox/, etc.

Comparison to Models Collection of data from various projects for a class of measurements is required if one wants to compare various chemical transport models or other models to the measured results available. Some well-known examples of such data collection activities are described by Emmons et al. [2000]. The “classical” approach for evaluating a model in such studies was to aggregate the observations over specific domains, altitude ranges, and time periods, and then to compare statistical quantities such as mean or median values and standard deviations for these aggregates with corresponding model data. However, the model fields were usually not sampled at exactly the same times and positions as the measurements but rather averages over entire time periods (e.g., monthly means) and domains (e.g., over a range of grid cells) were calculated because such quantities can easily be derived from standard model output. A more direct approach was used, e.g., by Brunner et al. [2003], compares each measured data point with its temporally and spatially interpolated model counterpart. Such a “point-by-point” analysis requires us to simulate exactly the same time periods when the measurements were obtained. Within the TRADEOFF project [Brunner et al., 2003], an extensive database including all major research aircraft and commercial airliner measurements between 1995 and 1998 as well as ozone soundings were established. The database is constructed to support direct comparison between model results and measurements. Moreover, quantitative methods were applied to judge model performance including the calculation of average concentration biases and the visualization of correlations and RMS errors in the form of so-called Taylor diagrams. The comparisons show quite good agreements for tracers that are sufficiently smooth (such as ozone) but often rather bad agreements for short-lived trace gases with ill-defined sources and sinks (such as NO). In particular, the NO measurements show a larger dynamic range than the models. The correlation coefficient between modeled and measured results, see Figure 6 for example, is often far below 0.5. This indicates that still a lot of development work and higher numerical model resolution is required to bring chemical transport models into a state of satisfying accuracy.

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1.000

Dec Jan Feb Mar

NOXAR NOx / ppbv

Apr May

0.100

Jun Jul Aug Aep Oct

0.010

r = 0.42

0.001 0.001

0.010

0.100

1.000

model NOx / ppbv

FIGURE 6. Scatter plot of model results versus NOXAR NO [Brunner et al., 2003].

Conclusions Research aircraft form an essential component of the atmospheric (and Earth observation) research infrastructure. Usage of aircraft requires skills, as indicated in this lecture, not only with respect to the underlying science, but also with respect to aviation, instrumentation, campaign management, and data evaluation.

References Baehr, J., H. Schlager, H. Ziereis, P. Stock, P. van Velthoven, R. Busen, J. Ström, U. Schumann, Aircraft observations of NO, NOy, CO, and O3 in the upper troposphere from 60 °N to 60 °S—Interhemispheric differences at mitlatitudes, Geophys. Res. Lett. 30(1), 10.1029/2003GL016935, 2003. Brenninkmeijer, C.A.M., Crutzen, P.J., Fischer, H., Güsten, H., Hans, W., Heinrich, G., Heintzenberg, J., Hermann, M., Immelmann. T., Kersting, D., Maiss, M., Nolle, M., Pitscheider, A., Pohlkamp, H., Scharffe, D., Specht, K., and Wiedensohler, A. (1999) CARIBIC—civil aircraft for global measurement of trace gases and aerosols in the tropopause region, J. Atmos. Oceanic Technol. 16, 1373–1383, 1999. Brough, N., C. Reeves, S.A. Penkett, K. Dewey, J. Kent, H. Barjat, P. S. Monks, H. Ziereis, P. Stock, H. Huntrieser and H. Schlager, Intercomparison of aircraft instruments on board the C-130 and Falcon 20 over southern Germany during EXPORT 2000, ACP, 2003. Brunner, B., J. Staehelin, D. Jeker, H. Wernli, and U. Schumann, Nitrogen oxides and ozone in the tropopause region of the Northern Hemisphere: Measurements from commercial aircraft in 1995/96 and 1997. J. Geophys. Res., 106, 27673–27699, 2001.

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Brunner, D., Staehlin, J., Rogers, H.L., Köhler, M., Pyle, J., Hauglustaine, L., Jourdain, T., Berntsen, K., Gauss, M., Isaksen, I., Meijer, E., Velthoven van, P., Pitari, G., Mancini, E., Grewe, V., Sausen, R., 2003: An evaluation of the performance of chemistry transport models by comparison with research aircraft observations. Part 1: Concepts and overall model performance. Atmos. Chem. Phys., 3, 1609–1631, 2003. Ehret, G., K.P. Hoinka, J. Stein, A. Fix, C. Kiemle, G. Poberaj, Low-stratospheric water vapor measured by an airborne DIAL. J. Geophys. Res., 104, D24, 31,351–31,359, 1999. Ehret, G., H.H. Klingenberg, U. Hefter, A. Assion, A. Fix, G.S. Poberaj, S. Berger, S. Geiger, Q. Lü, High peak and average power all-solid-state laser systems for airborne LIDAR applications. LaserOpto, 32(1), 29–37, 2000. Emmons, L. K., D.A. Hauglustaine, J.-F. Müller, M.A. Carroll, G.P. Brasseur, D. Brunner, J. Staehelin, V. Thouret, A. Marenco, Data composites of airborne observations of tropospheric ozone and its precursors, J. Geophys. Res., 105, (20), 497–538, 2000. Fix, A., M. Wirth, A. Meister, G. Ehret, M. Pesch, D. Weidauer, Tunable ultraviolett optical parametric oscillator for differential absorption Lidar measurements of tropospheric ozone. Appl. Phys. B, 75, (10.1007/s00340-002-0964-y), 153–163, 2002. Flatoy, F., O. Hov, H. Schlager, Chemical forecasts used for measurement flight planning during POLINAT 2. Geophys. Res. Lett., 27(7), 951–954, 2000. Gayet, J.-F., F. Auriol, A. Minikin, J. Strom, M. Seifert, R. Krejci, A. Petzold, G. Febvre and U. Schumann, Quantitative measurements of the microphysical and optical properties of cirrus clouds with four different in situ probes: Evidence of small ice crystals, Geophys. Res. Lett., 29(24), 2230, doi:10.1029/2001GL014342, 2002. Gayet, J.-F., J. Ovarlez, V. Shcherbakov, J. Ström, U. Schumann, A. Minikin, F. Auriol, A. Petzold, M. Monier, Cirrus cloud microphysical and optical properties at southern and northern midlatitudes during the INCA experiment. J. Geophys. Res., 109, D20206, doi:10.1029/ 2004JD004803, 2004. Heland, J., H. Schlager, A. Richter, J.P. Burrows, First comparison of tropospheric NO2 column densities retrieved from GOME and in situ aircraft profile measurements. Geophys. Res. Lett., 29, 20 (10.1029/2002GL015528), 44-1-44-4, 2002. Helten, M., H.G.J. Smit, D. Kley, J. Ovarlez, H. Schlager, R. Baumann, U. Schumann, P. Nedelec, A. Marenco, In-flight comparison of MOZAIC and POLINAT water vapor measurements. J. Geophys. Res., 104, D21, 1999, 26,087-26,096. Höller, H., U. Finke, H. Huntrieser, M. Hagen, and C. Feigl, Lightning produced NOX (LINOX)—Experimental design and case study results, J. Geophys. Res., 104, 13911–13922, 1999. Huntrieser, H., Ch. Feigl, H. Schlager, F. Schröder, Ch. Gerbig, P. van Velthoven, F. Flatøy, C. Théry, A. Petzold, H. Höller, U. Schumann, Airborne measurements of NOx, tracer species and small particles during the European Lightning Nitrogen Oxides Experiment, J. Geophys. Res., 107(11), AAC 2-1–AAC 2-27, 10.1029/2000JD000209, 2002. Minikin, A., A. Petzold, J. Ström, R. Krejci, M. Seifert, P. van Velthoven, H. Schlager, U. Schumann, Aircraft observations of the upper tropospheric fine particle aerosol in the Northern and Southern Hemispheres at midlatitudes, Geophys. Res. Lett. 30(10), 10.1029/2002GL016458, 2003. Ovarlez, J., J.-F. Gayet, K. Gierens, J. Ström, H. Ovarlez, F. Auriol, R. Busen, U. Schumann, Water vapour measurements inside cirrus clouds in Northern and Southern hemispheres during INCA, Geophys. Res. Lett., 29(16), 60-1–60-4, 200210. 1029/2001 GL014440, 2002. Schlager, H., P. Konopka, P. Schulte, U. Schumann, H. Ziereis, F. Arnold, M. Klemm, D., E. Hagen, P. D. Whitefield, J. Ovarlez, In situ observations of airtraffic emission signatures in the North Atlantic flight corridor. J. Geophys. Res., 102, 10,739–750, 1997.

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Schlager, H., P. Schulte, F. Flatoy, F. Slemr, P. van Velthoven, H. Ziereis, and U. Schumann, Regional nitric oxides enhancements in the North Atlantic flight corridor observed and modeled during POLINAT 2—a case study, Geophys. Res. Lett., 26, 3061–3064, 1996. Schumann, U., P. Konopka, R. Baumann, R. Busen, T. Gerz, H. Schlager, P. Schulte, H. Volkert: Estimate of diffusion parameters of aircraft exhaust plumes near the tropopause from nitric oxide and turbulence measurements. J. Geophys. Res., 100, 14147–14162, 1995. Schumann, U., H. Schlager, F. Arnold, J. Ovarlez, H. Kelder, Ø. Hov, G. Hayman, I. S. A. Isaksen, J. Staehelin, and P. D. Whitefield, Pollution from aircraft emissions in the North Atlantic flight corridor: Overview on the POLINAT projects, J. Geophys. Res., 105, 3605–3631, 2005. Schumann, U., F. Arnold, R. Busen, J. Curtius, B. Kärcher, A. Kiendler, A. Petzold, H. Schlager, F. Schröder, K.-H. Wohlfrom, Influence of fuel sulfur on the composition of aircraft exhaust plumes: The experiments SULFUR 1–7, J. Geophys. Res., 107(D15), 10.1029/2001JD000813, AAC 2-1 – AAC 2-27, 2002. Schumann, U., H. Huntrieser, H. Schlager, L. Bugliaro, C. Gatzen, H. Hoeller, Nitrogen Oxides from thunderstorms—Results from experiments over Europe and the Continental tropics, DACH—Deutsch-Österreichisch-Schweizerische MeteorologenTagung, 7.–10. September 2004, Karlsruhe, Deutschland, Proceedings in press (2004), http://imk-msa.fzk.de/dach2004/ Ström, J., M. Seifert, J. Ovarlez, A. Minikin, J.-F. Gayet, R. Krejci, A. Petzold, F. Auriol, R. Busen, U. Schumann, B. Kärcher, W. Haag, and H.-C. Hansson, Cirrus cloud occurrence as function of ambient relative humidity: A comparison of observations obtained during the INCA experiment, Atmos. Chem. Phys., 3, 1807–1816, 2003. Ziereis, H., H. Schlager, P. Schulte, I. Köhler, R. Marquardt, C. Feigl, In situ measurements of the NOx distribution and variability over the Eastern North Atlantic. J. Geophys. Res., 104, 16021–16032, 1999. Ziereis, H., H. Schlager, P. Schulte, P. E. J. van Velthoven, F. Slemr, Distributions of NO, NOx and NOy in the upper troposphere and lower stratosphere between 28° and 61° during POLINAT 2. J. Geophys. Res., 105, D3, 3653–3664, 2000. Ziereis, H., A. Minikin, H. Schlager, J. F. Gayet, F. Auriol, P. Stock, J. Baehr, A. Petzold, U. Schumann, A. Weinheimer, B. Ridley, J. Ström, Uptake of reactive nitrogen on cirrus cloud particles during INCA. Geophys. Res. Lett., 31(5), L05115 10.1029/2003GL018794, 2004.

Chapter 7 MOZAIC—Measuring Tropospheric Constituents from Commercial Aircraft JEAN-PIERRE CAMMAS

Introduction The overall objective of MOZAIC is to improve our knowledge on the concentrations and life cycles of ozone, water vapor, and related trace gases (CO and NOy) in the atmosphere. Particular emphasis is on the upper troposphere and lower stratosphere (UT/LS), a region where large gaps exist in our present knowledge and which is of great importance for climate. Three phases of MOZAIC have been funded by the European Commission from 1993 to February 2004. In the first phase of MOZAIC (1994–1996) automatic devices for measurements of ozone and water vapor have been installed on 5 Airbus A340. Lufthansa, Air France, Austrian, and Sabena fly the MOZAIC instrumentation free of charge. In the second phase of MOZAIC (1997–2000), O3 and H2O measurements were continued while new devices for measurements of carbon monoxide (CO) and nitrogen oxides (NOy) were built. In the third phase of MOZAIC (2001–2004), new devices for CO were installed onboard the five aircraft and the NOy instrument was installed onboard one aircraft. From August 1994 to December 2003, 22,003 flights (156,652 flights hours) have been performed by the aircrafts at a quasi-global scale. The MOZAIC consortium reflects a unique collaboration between scientists (CNRS France, Forschungszentrum Germany, Météo-France, University of Cambridge UK), the aircraft manufacturer Airbus, the airlines (Lufthansa, Air France, and Austrian) that transport free of charge the MOZAIC instrumentation, and the European Commission.

Instrumentation The MOZAIC ozone instrument is an improved version of a commercial dual-beam UV absorption instrument [Marenco et al., 1998; Thouret et al., 1998]. The accuracy is ±2 ppbv, the precision ±2%, and the response time 4 sec. The instrument is replaced every C-check of the aircraft (about 6 months) and is recalibrated at the laboratory.

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The MOZAIC relative humidity instrument is a capacitive RH sensor (Humicap-H) mounted in a Rosemount Probe System (Helten et al., 1998). The accuracy is less than 10%, the precision is about ±7%, the response time depends on temperature (a few secondes at ground to 2 minutes at cruise altitude). Sensor replacement and pre- and postcalibration are made in laboratory every C-check of the aircraft. The MOZAIC CO instrument is an improved version of a commercial IR gas filter correlation model. Its accuracy is ±5 ppbv, precision ±5%, detection limit 10 ppbv, response time 30 sec [Nédélec et al., 2003]. Sensor replacement and preand postcalibration are made in laboratory every C-check of aicraft. The aeronautical certification for the instruments above has been delivered by Airbus. The MOZAIC NOy instrument detects NO by chemiluminescence with O3 in combination with catalytic conversion of the other NOy compounds to NO at 300 °C on a gold surface in the presence of H2 [Volz-Thomas et al., 2004]. Sensitivity of 0.3–0.5 cps/ppt gives a detection limit of better than 30–50 ppt for an integration time of 4 s, and 150–300 ppt at the maximum resolution of the instrument (10 Hz). It is designed for unattended operation during 400–800 flight hours. The total weight is 50 kg, including calibration system, compressed gases, mounting, and safety measures. Aeronautical certification is delivered by Lufthansa technics.

Database The MOZAIC database is located in Toulouse at CNRM (Centre National de la Recherche Météorologique, Météo-France). It contains raw data (full resolution) and averaged data (lower resolution for the purpose of comparison with results from large-scale models). It also contains other relative meteorological data, like the difference of pressure between the aircraft and the dynamical tropopause on ECMWF analyses (taken at 2 units of potential vorticity), or like quick views of satellite images. The access to MOZAIC data is public and subject to the signature of a protocol by co-investigators. Details of the protocol are provided at the end. About 50 international groups are using MOZAIC data.

Main Results Scientific achievements: 1. Climatology of ozone at a quasi-global scale (a) in the UTLS at mid-latitudes showing (i) a spring maximum in the LS and a spring-to-summer maximum in the UT, (ii) a large interannual variability with a strong positive anomaly

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3.

4.

5.

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in 1998–1999 from Northeast America to Europe, (b) on vertical profiles showing (i) the 1998–1999 anomaly over Paris, Frankfurt, and New York, (ii) ozone maxima confined in the lower troposphere over Africa during the dry season in relation to biomass burning and photochemistry. Climatology of relative humidity showing that the UT is considerably moister than previously assumed. Over the North Atlantic, more than 30% of the data show substantial ice super-saturation (ISS). This finding is not reflected in the ECMWF analyses, which significantly underestimate the measurements above a relative humidity of 50%. MOZAIC shows that ISS is a small-scale phenomenon and that the distribution of humidity in the tropical UT is bi-modal. First quasi-global scale distributions of CO and NOy in the UTLS region showing the strong impact of small-scale convection and large-scale uplifting by warm conveyor belts on the composition of the UT, especially over the east coast, of North America and over Asia. Quantitative assessments of stratosphere-troposphere exchanges (subtropical filaments associated with Rossby wave breaking, mid-latitude tropopause folds, …) using a variety of Lagrangian methodologies applied to numerical weather analyses and using mesoscale and global models. Assessment of present modeling capabilities for the UTLS region by Chemistry Transport Models through validation studies against MOZAIC data, showing a relatively good ability of models tested (TOMCAT, Cambridge; MOCAGE, Météo-France, MOZART-2, MPI-Hamburg) to reproduce the ozone cycle in the UTLS and identifying some deficiencies in the vertical transport and in the parameterizations of clouds and chemistry. Numerical assessments by these models of the perturbation of the UTLS region due to aircraft emissions fall within the most recent values published in the literature. Detection of recent aircraft emissions with 10-Hz measurements of the NOy device. A climatological survey shows that the such plumes comprise about 1.5% of the total NOy over the North Atlantic flight corridor, USA, and Europe, without significant spatial or seasonal variations. Smaller fractions are observed over North Africa and much higher fractions occur over East Asia, most likely because of narrower corridors in this region.

Main deliverables: 1. Ten years of quality-assured measurements of ozone (O3) and water vapor (H2O); more than 2 years of quality-assured measurements of carbon monoxide (CO), 3 years of total odd-nitrogen (NOy); from 5 commercial aircraft between 0 and 12 km altitude over the period 1994–2004 (> 22,000 intercontinental flights). The database is open to public access. 2. Automatic instruments with operating procedures and data retrieval algorithms for routine measurements of O3, H2O, CO, and NOy aboard in-service aircraft. 3. Improved knowledge of the climatology and interannual variability of O3 in the UTLS region.

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4. Improved knowledge of the distribution of water vapor in the UT, on the quality of its representation by ECMWF analyses, on the ice super-saturation phenomena, and on processes responsible for its distribution in the tropical UT. 5. First large-scale distributions of CO and NOy in the UTLS region, including seasonality, and of vertical profiles at quasi-global scale. 6. Improved knowledge of the impact of (i) stratospheric intrusions on the composition of the troposphere, and (ii) of vertical transport from the boundary layer on the chemical composition of the UT region. 7. Improvement of our capacity in validating Chemistry Transport Models, which are the tools used for quantification of the UTLS ozone budget, for assessment of climate change and impact of aircraft, and for providing answers to policy makers and industry.

Future of Mozaic A memorandum of understanding has been signed between Airbus, Lufthansa, Air France, Austrian, and Centre National de la Recherche Scientifique to continue the present MOZAIC phase until December 2007. The internationally expressed need for representative high-quality in-situ data for the UTLS and the technical expertise developed in MOZAIC lead to an initiative for extending the MOZAIC concept into a large European infrastructure that will provide real-time in-situ measurements of trace gases in the UTLS region to the European GMES initiative. As a first step, a design study for new infrastructures (IAGOS: Integration of routine aircraft measurements into the Global Observing System) submitted in the last call of FP6 for the design and certification of new instrumentation has been accepted by the European Commission and has started in April 2005 for a 3-year duration.

Data Access Protocol Each MOZAIC co-investigator is asked to sign the MOZAIC data protocol (http://www.aero.obs-mip.fr/mozaic/) and to provide piece of information on: ●

●

The research that will be performed in the next two years with use of MOZAIC data (1 page maximum), The MOZAIC Principal Investigator(s) to which collaboration or/and coauthorship could be proposed.

Principal investigators of the MOZAIC program: J.-P. Cammas (LA/CNRS, Toulouse, France, MOZAIC-III coordinator, [email protected]) V. Thouret (LA/CNRS, Toulouse, France, [email protected]) H. Smit (FZJ, Jüliech, Germany, [email protected]) F. Karcher (Météo-France, Toulouse, France, [email protected]) A. Volz-Thomas (FZJ, Jüliech, Germany, [email protected])

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References Asman, W.A.H., M.G. Lawrence, C.A.M. Brenninkmeijer, P.J. Crutzen, J.W.M. Cuijpers, P. Nédélec, Rarity of upper-tropospheric low O3 mixing ratio events during MOZAIC flights. Atmos. Chem. Phys. 3, 1541–1549, 2003. Baray J.L., T. Randriambelo, S. Baldy, G. Ancellet, Comment on ‘Tropospheric O3 distribution over the Indian Ocean during spring 1995 evaluated with a chemistry-climate model’, by A.T.J. de Laat et al., Journal of Geophysical Research, 106(D1), 1365–1368, 2001. Baray J.L., S. Baldy, R.D., Diab, J.P. Cammas, Dynamical study of a tropical cut-off low over South Africa, and its impact on tropospheric ozone, Atmospheric Environment, 37(11), 1475–1488, 2003. Borchi F., A. Marenco, Discrimination of air masses near the extratropical tropopause by multivariate analyses from MOZAIC data. Atmos. Environ. 7, 1123–1135, Mar. 2002. Bregman A., M.C. Krol, H. Teyssedre, W.A. Norton, A. Iwi, M. Chipperfield, G. Pitari, J.K. Sundet, J. Lelieveld, Chemistry-transport model comparison with ozone observations in the midlatitude lowermost stratosphere. J. of Geophys. Res., 106(D15), 17479–17496, 2001. Bregman, B., A. Segers, M. Krol, E. Meijer and P. van Velthoven: On the use of massconserving wind fields in chemistry-transport models. Atmos. Chem. Phys. 3, 447–457, 2003. Brioude et al., MOZAIC case study of stratosphere-troposphere exchanges in a summer extratropical low, To be submitted to Atmos. Chem. Phys Discuss., 2005. Cammas J.-P., S. Jakobi-Koali, K. Suhre, R. Rosset, A. Marenco, Atlantic subtropical potential vorticity barrier as seen by MOZAIC flights. J. of Geophys. Res., 103, 25681–25693, 1998. Cathala, M.L., J. Pailleux, et V.-H. Peuch, Improving global simulations of UTLS ozone with assimilation of MOZAIC data. Tellus, 55B, 1–10, 2003. Cho, J.Y.N., V. Thouret, R.E. Newell, A. Marenco, Isentropic scaling analysis of ozone in the upper troposphere and lower stratosphere, J. Geophys. Res., 106, 10,023–10,038, 2000. Cho, J.Y.N., E. Lindborg, Horizontal velocity structure functions in the upper troposphere and lower stratosphere: 1. Observations, J. Geophys. Res., 106, 10,223–10,232, 2001. Cooper O.R., A. Stohl, S. Eckhardt, D. Parrish, S.J. Oltmans, B.J. Johnson, P. Nedelec, F.J. Schmidlin: A comparison of the east and west coasts of the United States during spring 2002: Different pollutant transport pathways but nearly identical tropospheric ozone burden. Submitted to JGR. Crowther, R.A., K.S. Law, J.A. Pyle, S. Bekki, H.G.J. Smit, Characterising the effect of large-scale model resolution upon calculated OH production using MOZAIC data, Geophys. Res. Lett., DOI 10.1029/2002GL014660, 2002. Crowther, R.A., K.S. Law, J.A. Pyle, A. Volz-Thomas, W. Paetz, P. Nedelec, J.P. Cammas, Investigating the origin of O3 in the upper troposphere using MOZAIC data, to be submitted to Atmos. Chem. Phys., 2005. Crowther, R.A., K.S. Law, J.A. Pyle, H.G.J. Smit, P. Nedelec, Evaluation of modelled photochemical OH production using MOZAIC data, to be submitted to Atmos. Chem. Phys., 2005. Dethof A, A. O’Neill, J.M. Slingo, et al., A mechanism for moistening the lower stratosphere involving the Asian summer monsoon. Quart. J. of the Roy. Met. Soc., 125(556): 1079–1106, 1999. Diab R., A. Raghunandan, A. Thompson, V. Thouret, Classification of tropospheric ozone profiles based on MOZAIC aircraft data. Atmos. Chem. and Phys., 3, 713–723, 12-6-2003.

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Edwards, D.P., J.-F. Lamarque, J.-L. Atti\’e, L.K. Emmons, A. Richter, J.-P. Cammas, J.C. Gille, G.L. Francis, M.N. Deeter, J. Warner, D. Ziskin, L.V. Lyjak, J. R. Drummond, J. P. Burrows, Tropospheric ozone over the tropical Atlantic: A satellite perspective, J. Geophys. Res. 108, 4237, doi:10.1029/2002JD002927, 2003. Emmons, L.K., D.A. Hauglustaine, J-F. Muller, M.A. Caroll, G.P. Brasseur, D. Brunner, J. Staehelin, V. Thouret, A. Marenco, Data composites of airborne observations of tropospheric ozone and its precursors, J. Geophys. Res., 105, 20,497–20,538, 2000. Forster, C., A. Stohl, P. James, V. Thouret, The residence times of aircraft emissions in the stratosphere using a mean emissions inventory and emissions along actual flight tracks. J. Geophys. Res., doi 10.1029/D12 2002JD0002515 8524, 2003. Gauss M, G. Myhre, G. Pitari, M.J. Prather, I.S.A. Isaksen, T.K. Berntsen, Brasseur, G.P. Dentener FJ, Derwent RG, Hauglustaine DA, Horowitz LW, Jacob DJ, Johnson M, Law KS, Mickley LJ, Muller JF, Plantevin PH, Pyle JA, Rogers HL, Stevenson DS, Sundet JK, van Weele M, Wild O, Radiative forcing in the 21st century due to ozone changes in the troposphere and the lower stratosphere. J. Geophys. Res., 108(D9): Art. No. 4292 May 13, 2003. Giannakopoulos, C., P. Good, P.K.S. Law, D.F. Shallcross and K.Y. Wang, Modelling the impacts of aircraft traffic on the chemical composition of the upper troposphere, Proc. Inst. Mech. Eng., J. Aerosp. Eng., 217(G5), 237–243, 2003. Gheusi, F., J.-P. Cammas, F. Cousin, C. Mari, P. Mascart, Quantification of mesoscale transport across the frontiers of the free troposphere: a new method and applications to ozone. Atmos. Chem. Phys. Discussion, Sept. 2004. Gierens, K.M., U. Schumann, H.G.J. Smit, M. Helten, G. Zängl, Determination of humidity and temperature fluctuations based on MOZAIC data and parametrization of persistent contrail coverage for general circulation models. Ann. Geophys. 15, 1057–1066, 1997. Gierens, K., U. Schumann, M. Helten, H.G.J. Smit, A. Marenco: A distribution law for relative humidity in the upper troposphere and lower stratosphere derived from three years of MOZAIC measurements. Ann. Geophys. 17, 1218–1226, 1999. Gierens, K., P. Spichtinger: On the size distribution of ice-supersaturated regions in the upper troposphere and lowermost stratosphere. Ann. Geophys. 18, 499–504, 2000. Gierens, K., U. Schumann, M. Helten, H. Smit, P.H. Wang, Ice-supersaturated regions and subvisible cirrus in the northern midlatitude upper troposphere. J. Geophys. Res. 105, 22743–22754, 2000. Gierens, K., R. Kohlhepp, P. Spichtinger, M. Schroedter-Homscheidt, Ice supersaturation as seen from TOVS. Atmos. Chem. Phys. Discuss., 4, 299–325, (http://www.copernicus. org/EGU/acp/acpd/299/acp-4-299_p.pdf) Gouget H., G. Vaughan, A. Marenco, H. G. J. Smit., Decay of a cut-off low and contribution to stratosphere-troposphere exchange. Quart. J. Roy. Met. Soc., 126, 1117–1142, 2000. Helten M., H., Smit, W. Strater, D., Kley, P., Nedelec, M., Zoger, R. Busen, Calibration and performance of automatic compact instrumentation for the measurement of relative humidity from passenger aircraft, J. Geophys. Res., 103, 25,643–25,652, 1998. Helten M., H.G.J. Smit, D. Kley, J. Ovarlez, H. Schlager, R. Baumann, U. Schumann, P. Nedelec, A. Marenco, In-Fight Intercomparison of MOZAIC and POLINAT water vapor measurements, J. Geophys. Res., 104, D21, 26,087–26,096, 1999. Law, K.S., P.-H. Plantevin, D.E. Shallcross, H. Rogers, C. Grouhel, V. Thouret, A. Marenco, J.A. Pyle, Evaluation of modelled O3 using MOZAIC data, J. Geophys. Res., 103, 25721–25740, 1998. Law, K. S., P.H. Plantevin, V. Thouret, A. Marenco, W. Asman, W.A.H. Lawrence, P.J. Crutzen, J.F. Muller, D.A. Hauglustaine, M. Kanakidou, Comparison between global

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chemistry transport model results and measurements of ozone and water vapour by airbus in-service aircraft (MOZAIC) data, J. Geophys. Res., 105, 1503–1525, 2000. Lee, S.H., M. Le Disloquer, R. Singh, S.E. Hobbs, C. Giannakopoulos, P.H. Plantevin, K.S. Law, J.A. Pyle and M.J. Rycroft, Implication of NOy emissions from subsonic aircraft at cruise altitude, Proc. Inst. Mech. Eng. J. Aerosp. Eng., 211, G3, 157–168, 1997. Li, Q., D. J. Jacob, J. A. Logan, I. Bey, R. M. Yantosca, H. Liu, R. V. Martin, A. M. Fiore, B. D. Field, B. N. Duncan, V. Thouret, A tropospheric ozone maximum over the Middle East, Geophys. Res. Lett., 28, 3235–3238, 2001. Lindborg E., Can the atmospheric kinetic energy spectrum be explained by two-dimensional turbulence?, J. Fluid. Mech., 388, 259–288, 1999. Lindborg, E., J. Y.N. Cho, Determining the cascade of passive scalar variance in the lower stratosphere, Phys. Rev. Lett., 85, 5663–5666, 2000. Lindborg E. and Cho J. Y.N., Horizontal velocity structure functions in the upper troposphere and lower stratosphere: 2. Theoretical considerations, J. Geophys. Res., 106 (D10), 10233–10241, 2001. Marenco, A., V. Thouret, P. Nedelec, H. Smit, M. Helten, D. Kley, F. Karcher, P. Simon, K. Law, J. Pyle, G. Poschmann, R. Von Wrede, C. Hume, T. Cook, Measurement of ozone and water vapour by Airbus in-service aircraft: The MOZAIC airborne program, An overview, J. Geophys. Res., 103, 25631–25642, 1998. Martin, R.V., D.J. Jacob, J.A. Logan, I. Bey, R.M. Yantosca, A.C. Staudt, Q. Li, A.M. Fiore, B.N. Duncan, H. Liu, P. Ginoux, V. Thouret, Global model analysis of TOMS and in-situ observations of tropical tropospheric ozone, J. Geophys. Res., 107(D18), 4351, doi: 10.1029/2001JD001480, 2002. Morgenstern, O., A. Marenco, Wintertime climatology of MOZAIC ozone based on the potential vorticity and ozone analogy. J. Geophys. Res., 105(D12), 15481–15493, 2000. Morgenstern, O., G.D. Carver, Comparison of cross-tropopause transport and ozone in the upper troposphere and lower stratosphere region, J Geophys Res-Atmos 106(D10): 10205–10221 May 27, 2001. Nedelec P., J.-P. Cammas, V. Thouret, G. Athier, J.-M. Cousin, C. Legrand, C. Abonnel, F. Lecoeur, G. Cayez, C. Marizy, An improved infra-red carbon monoxide analyser for routine measurements aboard commercial Airbus aircraft: Technical validation and first scientific results of the MOZAIC III Program. Submitted to Atmos. Chem. And Phys., 3, 1551–1564, 29-9-2003. Nedelec P., V. Thouret, J. Brioude, B. Sauvage, J.-P. Cammas, A. Stohl, Extreme CO concentrations in the upper troposphere over North-East Asia in June 2003 from the in-situ MOZAIC aircraft data. Geophys. R. Letters, accepted, 2005. Nedoluha, G.E., R.M. Bevilacque, K.W. Hoppel, J.D. Lumpe, H.G.J. Smit, POAM III measurements of water vapor in the upper troposphere and lower most stratosphere, J. Geophys. Res., 107, 10.1029/2001JD000793, 2002. Newell R., V. Thouret, J. Cho, P. Stoller, A. Marenco, H. Smit, Ubiquity of quasi-horizontal layers in the troposphere, Nature, 398, 316–319, 1999. O’Connor, F.M., K. S. Law, J. A. Pyle, et al., Tropospheric ozone budget: regional and global calculations, Atmos. Chem. Phys. Discussions, MS-NR: acp2004–001, 2004. Offermann D., B. Schaeler, M. Riese, M. Langfermann, M. Jarisch, G. Eidmann, C. Schiller, H.G.J. Smit, W.G. Read, Water vapor at the tropopause during the CRISTA 2 mission, J. Geophys. Res., 107, DOI 10, 1029/2001JD000700, 2002. Prados A.I., G.E. Nedoluha, R.M. Bevilacqua, D.R. Allen, K.W. Hoppel, A. Marenco, POAM III ozone in the upper troposphere and lowermost stratosphere: Seasonal

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variability and comparisons to aircraft observations. J. Geophys. Res., 108, (D7), 10.1029/2002JD002819, 2003. Prather, M., M. Gauss, T. Berntsen, I. Isaksen, J. Sundet, I. Bey, G. Brasseur, F. Dentener, R. Derwent, D. Stevenson, L. Grenfell, D. Hauglustaine, L. Horowitz, D. Jacob, L. Mickley, M. Lawrence, R. von Kuhlmann, J.-F. Muller, G. Pitari, H. Rogers, M. Johnson, J. Pyle, K. Law, M. van Weele, O. Wild, Fresh air in the 21st century? Geophys. Res. Letts., 30(2), 1100, doi:10.1029/2002GL016285, 2003. Ravetta F., G. Ancellet. Identification of dynamical processes at the tropopause during the decay of a cut-off low using high resolution airborne lidar ozone measurements. Mon. Weather Rev., 128, 3252–3267, 2000. Roelofs G.J., A.S. Kentarchos, T. Trickl, A. Stohl, W.J. Collins, R.A. Crowther, D. Hauglustaine, A. Klonecki, K.S. Law, M.G. Lawrence, R. von Kuhlmann, M. van Weele, Intercomparison of tropospheric ozone models: Ozone transport in a complex tropopause folding event, 108(D12): art. no.8529, J. Geophys. Res., 2003. Sauvage et al., Tropospheric ozone over Equatorial Africa: regional aspects from the MOZAIC data. Atmos. Chem. Phys., 4, 3285–3322, 2005. Scott, R.K., J.-P., Cammas, P. Mascart et al., Stratospheric filamentation into the upper tropical troposphere. J. Geophys. Res., 106(D11), 11835–11848, 2001. Scott, R.K., J.-P. Cammas, Wave breaking and mixing at the subtropical tropopause. J. of Atmos. Sciences. 59(15), 2347–2361, 2002. Scott, R.K., E. Schuckburgh, J.-P. Cammas, B. Legras, Stretching rates and equivalent length near the tropopause. J. of Geophys. Res., D13 2002JD002988, 4394. Spichtinger P., K. Gierens, W. Read, The statistical distribution law of relative humidity in the global tropopause region. Meteorol Z, 11(2), 83–88, 2002. Spichtinger, P., K. Gierens, U. Leiterer, H. Dier, Ice supersaturated region over Lindenberg, Germany. Meteorol. Z., 12, 143–156, 2003. Spichtinger, P., K. Gierens, W. Read, The global distribution of ice-supersaturated regions as seen by the Microwave Limb Sounder. Q. J. R. Meteorol. Soc., 129, 3391–3410, 2003. Stohl, A., T. Trickl, A textbook example of long-range transport: Simultaneous observation of ozone maxima of stratospheric and North American origin in the free troposphere over Europe. J. Geophys. Res., 104, 30445–30462, 1999. Stohl, A., P. James, C. Forster, N. Spichtinger, A. Marenco, V. Thouret, H. G. J. Smit, An extension of measurements of ozone and water vapor by Airbus in-service aircraft (MOZAIC) ozone climatologies using trajectory statistics. J. Geophys. Res. 106, 27,757–27,768, 2001. Suhre K., J.-P. Cammas, P. Nedelec, R. Rosset, A. Marenco, H.G.J. Smit, Observations of high ozone transients in the upper Equatorial Atlantic troposphere. Nature, 388, 661–663, Aug. 1997. Spichtinger, P., K. Gierens, H.G.J. Smit, J. Ovarlez, J.-F. Gayet, On the distribution of relative humidity in cirrus clouds. Atmos. Chem. Phys. Discuss., 4, 365–397, 2004. (http:// www.copernicus.org/EGU/acp/acpd/365/acpd-4-365_p.pdf). Thouret, V., A. Marenco, J. A. Logan, P. Nedelec, C. Grouhel, Comparisons of ozone measurements from the MOZAIC airborne program and the ozone sounding network at eight locations, J. Geophys. Res., 103, 25,695–25,720, 1998. Thouret, V., A. Marenco, P. Nédélec, C. Grouhel, Ozone climatologies at 9–12 km altitude as seen by the MOZAIC airborne program between September 1994 and August 1996, J. Geophys. Res., 103, 25,653–25,679, 1998. Thouret V.J.Y.N. Cho, R.E. Newell, et al., General characteristics of tropospheric trace constituent layers observed in the MOZAIC program. J. Geophys. Res., 105(D13), 17379–17392, 2000.

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Thouret V, J.Y.N. Cho, M.J. Evans, et al., Tropospheric ozone layers observed during PEM-Tropics B. J. Geophys. Res., 106(D23), 32527–32538, 2001. Thouret et al., Tropopause-referenced ozone climatology and interannual variability (1994–2000) from the MOZAIC program. Submitted to Atmos. Chem. Phys. Discuss, 2005. Tulet P, K. Suhre, C. Mari, et al., Mixing of boundary layer and upper tropospheric ozone during a deep convective event over Western Europe. Atmos. Environ., 36(28), 4491– 4501, 2002. Zahn, A., C.A.M. Brenninkmeijer, W.A.H. Asman, P.J. Crutzen, G. Heinrich, H. Fischer, J.W.M. Cuijpers, P.F.J. van Veldhoven, The budget of O3 and CO in the upper troposphere: The CARIBIC aircraft results 1997–2001. J. Geophys. Res. 107(D17), art. no. 4337, 2003. Zbinden R., J.-P. Cammas, V. Thouet, P. Nédélec, F. Karcher, P. Simon, Midlatitude tropospheric ozone columns from the MOZAIC programme: climatology and interannual variability. Submitted to Atmos. Chem. And Physics Discussion, 2005.

Chapter 8 Uninhabited Aerial Vehicles: Current and Future Use PAUL A. NEWMAN

Introduction Uninhabited aerial vehicles (UAVs) have recently become available for scientific research. UAV development has largely been underwritten by the military for reconnaissance. UAVs such as the Global Hawk have demonstrated a highaltitude (>16 km), long-range (>2000 km), and heavy payload (<1000 kg) capability. These attributes make UAVs very attractive for Earth science research. A UAV is classified as a powered vehicle that does not carry a human operator, uses aerodynamic forces to provide vehicle lift, can fly autonomously or be piloted remotely, can be expendable or recoverable, and carries a payload. Remotely piloted vehicles (RPVs) are a subclass of UAVs that require a ground pilot to control the aircraft. Radio-controlled aircraft are not UAVs if they do not carry a payload. Satellites are not UAVs, since they do not use aerodynamic forces. On the other hand, satellites bear many similarities to UAVs, as will be discussed later. In this paper, we will first discuss the usage of conventional aircraft for scientific missions, and how conventional aircraft relate to both satellites and UAVs. In the third section we will discuss some specific examples of UAVs, give a recent history of UAV usages for Earth science and applications, and give a brief description of UAV operations. The fourth section will discuss advantages and disadvantages of UAVs, and the fifth section will discuss future directions for UAV usage.

Aircraft How Are Aircraft Used for Earth Observations? Earth scientists have extensively used aircraft, and their use continues to grow. Currently, aircraft are used for both remote sensing and in-situ sampling of the atmosphere. Aircraft are used for (1) field experiments such as the Cirrus Regional

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Study of Tropical Anvils and Cirrus Layers—Florida Area Cirrus Experiment (CRYSTAL-FACE), (2) monitoring and applications (e.g., forest fire overflights and reconnaissance), (3) satellite validation (e.g., SOLVE II), (4) instrument and algorithm development, and in the future, (5) as suborbital systems that are integrated into satellite programs. Aircraft are used to measure long-term events that have time scales of months to years, short-term events with time scales of days to weeks, and mesoscale events with time scales of minutes to hours. Examples of long-term events are the Antarctic ozone hole, pollution systems, and climate change. These events are continuous or typically recur on a regular basis. Long-term events have large spatial scales and are easily planned for. Short-term events are typically synoptic scale (1000–5000 km). Streamers of pollution or synoptic scale weather systems are classic examples. While these short-term events are not precisely predictable, they tend to regularly recur, and can be planned for. They are typically more difficult to sample than the long-term, large-scale events. The mesoscale systems are the most difficult to sample. A convective system with a spatial scale of 10 km and a time scale of 3–4 hours is extremely difficult to sample. The transit time from the airfield to the phenomena may see the entire development and collapse of these systems. The smallerscale phenomena with chaotic behavior are the most difficult phenomena to sample. In any case, all of these phenomena require clever planning and good forecast tools. The CRYSTAL-FACE (CF) mission provides an interesting example of how aircraft are used to sample phenomena. CF was a measurement campaign for investigating tropical cirrus cloud properties and formation. CF was conducted in July 2002 from Key West, Florida, and included a mix of aircraft, ground-based observations, radiosondes, and satellite instruments. CF used 6 aircraft: the NASA ER-2, the NASA WB-57F, the Scaled Composites Proteus, the University of North Dakota’s Citation, the CIRPAS Twin Otter, and the Naval Research Lab’s P-3 Orion. The CF location was chosen based upon the regular early afternoon daily convection that occurs in South Florida over the moist Everglades. This convection was driven by the hot and moist boundary that was triggered by the daily sea breeze. This convection detrained cirrus at about 13 km. A surface radar in the Everglades tracked the development of this convection and directed the aircraft to sample the evolving cirrus as it detrained. The ER-2 payload (Figure 1) was principally composed of remote sensing instruments, while the WB-57F (Figure 2) payload was composed of in-situ instruments. The ER-2 flew above the detraining cirrus at an altitude of 20 km, while the WB-57F typically flew in the cirrus. The ER-2 and WB-57F were flown together along the same track (ER-2 above and WB-57F below), such that they were sampling the same cirrus simultaneously. The ER-2 and WB-57F were also flown to validate the NASA Aqua, Terra, and TRMM satellite instruments. Since the ER-2 duplicated the capabilities of these satellites (see satellites in Figure 1), it was possible to both validate satellite remote sensing instruments directly, and to bootstrap the in-situ observations to the satellite observations using a forward radiative transfer model. This was

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Radiation Meas. system (Rams) Valero (Scripps)

Conical Scanning Sub-mm wave Imaging Radiometer (CoSSIR), Wang (GSFC), Evans (CU)

AMSU Microwave Temp. Profiler (MTP) Mahoney (JPL)

Meteorological Meas. System (MMS) Bui (Ames)

TRMM ER-2 Doppler Radar (EDOP)

Cloud Radar System (CRS) Heymsfield (GSFC), Mace (Utah)

Solar Spectral Flux Radiometers (SSFR) Pilewskie (Ames)

CloudSAT

CERES Vaisala Dropsonde Halveroson (UMBC), Starr (GSFC)

King, Platnick (GSFC)

Cloud Physics Lidar (CPL) Mc Gill (GSFC) CALIPSO

JPL Laser Hygrometer (JLH) Herman (JPL)

Aqua & Terra

FIGURE 1. NASA ER-2 payload diagram for the July 2002 CRYSTAL-FACE mission. Total payload was 2183 lbs (992 kg).

accomplished by directly comparing the in-situ observations (e.g., ice water content) to the ER-2 measurements, and then to the same measurement made by the satellite. The CF mission demonstrated key aspects of aircraft field missions. First, it was a clever choice of location for the investigation of cirrus detrainment. Second, it L. Wings Hatches Frostpoint, VIPS, CIN-100

Upper Equip Bay RAMS (Zenith), SSR

R. Wing Hatches CPI, JLH, SPP-100

L. Wings Pod CPS, MASP, MTP, Harvard H2O Cockpit Video camera

Nose MMS, PALMS, PT

R. Wing pod ALIAS, CLH Pallet 2 Pallet 1 Pallet 3 Aft Transition Pallet 4 Harvard RAMS (Nadir) NO/NOy CIMS, Argus, FCAS, CH4, O3 MACS, NMASS Tot. H2O, CO2 PARTHER

FIGURE 2. NASA WB-57F payload diagram for the July 2002 CRYSTAL-FACE mission. Total payload was approximately 4000 lbs (1820 kg).

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showed the need for excellent forecasting and nowcasting for directing flight operations. Third, they optimally used the platforms such that each platform’s data could be related to the other platforms. Fourth, they combined ground, aircraft, and satellite observations into a unified experiment. Finally, CF clearly demonstrated how aircraft could be operationally integrated with satellite systems. The challenge for UAVs is to both duplicate some of our current capabilities with manned aircraft, and to extend these capabilities.

How Do Aircraft Relate to Other Satellites? Satellites and aircraft have different spatial coverage, temporal coverage, resolution, lifetimes, and payloads. Earth sciences satellites are typically in low Earth orbits (LEO) or geosynchronous Earth orbits (GEO). The LEO orbits of satellites such as Aura or Terra give approximately twice per day global coverage using scanning type instruments with day/night viewing capability. For a specific target such as a cloud or ground site, the LEO satellite provides 2 views per day. A GEO orbiting satellites such as GOES-8 provides continuous coverage for about 1/6th of the Earth’s surface. In contrast, an aircraft can provide approximately 6–8 hours of continuous observations of a specific location. LEO orbits are relatively low (e.g., Aura orbits at an altitude of about 780 km); GEO orbits are very high (36,000 km). For trace gas observations, aircraft can provide in-situ observations with horizontal/vertical resolutions of 0.1/0.01 km; GEO instruments generally provide horizontal/vertical resolutions of 1/5 km, while LEO instruments have resolutions of 10/1 km. Aircraft make sporadic observations depending on campaigns, while satellites provide near-continuous observations over 2–10-year periods. Satellite payloads are typically a few kg to 2000 kg, and aircraft payloads are typically a few hundred kg to 15,000 kg for aircraft such as the NASA DC-8. Satellites and aircraft provide somewhat different capabilities. However, as was shown in the CF mission example, these platforms are mutually dependent. Because of their range and duration, UAVs begin to approach the capabilities of satellites but with a capability of providing much more detailed observations of specific phenomena.

How Do Manned Aircraft Relate to UAVs? Manned aircraft and uninhabited aerial vehicles (UAVs) have many similarities, but also have significant differences. The advantages of manned aircraft over a UAV are their (1) long heritages, (2) wide varieties, (3) immense payload, power, and volume capacities, (4) airspace access, (5) real-time command and control, and (6) ability to carry passengers and instrument operators. The drawbacks for manned aircraft are (1) short duration capabilities (≤ 10 hours), (2) modest range (< 3700 km), (3) inability to fly in dangerous situations, and (4) inability to fly in dirty environments (ash, nuclear debris, etc.). UAVs provide opportunities to overcome some of the shortcomings of manned aircraft. First, some UAVs are capable of very long-duration and long-range

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flights (> 5,000 km, > 24 hours). Second, some UAVs are expendable. Third, UAVs can be flown in dirty conditions where a pilot could be harmed. UAVs are capable of doing dumb, dangerous, and dirty missions. In particular, because of the shorter durations for manned aircraft, manned aircraft are best suited for reconnaissance (i.e., short snap-shot views of phenomena). Because of its longer endurance, the UAVs expand our capabilities to perform surveillance (i.e., continuous observations of phenomena). Figure 3 is a plot of aircraft range and altitude for a number of UAVs and manned aircraft. Also included in Figure 3 are the horizontal and vertical dimensions of various atmospheric phenomena. Some UAVs (e.g., the Global Hawk) significantly expand the envelope beyond our current manned aircraft capability.

Uninhabited Aerial Vehicles (UAVs) Examples There are dozens of UAVs that have been developed or are under development today (see: http://www.uavforum.com/vehicles/capabilities.htm or http://www.aiaa. org/images/PDF/WilsonChart.pdf). These UAVs have been mainly developed for military purposes. Two broad classes of UAVs are high-altitude, long-endurance 30

Helios

Performance Envelope

ER2 WB57

Altitude (km)

20

ALTAIR

AL-II

0

Synoptic Scale

0

ITCZ Width

Mesoscale

P3

Walker Circulation

DC8

AL-I

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H

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L 10

Max. polar O3 loss

GH

Pr

8000

4000

Tropical Tropopause SVC Convective Detrainment

Polar Tropopause Troposphere

Boundary Layer 12000

Range (nm)

FIGURE 3. Altitude versus range plot for manned aircraft and UAVs (in boxes). The aircraft payload weight is proportional to font size (truncated at 2000 lb and 600 lb), and the bold-faced font indicates payloads greater than 2000 lb. The thick line indicates the performance envelope for manned aircraft. The L indicates a Learjet, while the H indicates the NSF Hiaper Gulfstream V aircraft. The scales of various atmospheric phenomena are indicated for perspective.

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(HALE) UAVs and medium-altitude, long-endurance (MALE) UAVs. In general, a long-endurance is defined to be greater than about 10 hours, while a high-altitude is defined to be about about 50,000 feet (15.2 km). A number of UAVs are now used or have been used for Earth Sciences and applications. Table 1 shows a brief listing of UAVs that have been or have planned uses for Earth science research. The HALE Global Hawk has some of the most desirable characteristics: a long endurance, high altitude, and long-range capability. The Global Hawk is powered by a conventional turbofan jet engine. In contrast, the Helios and the Pathfinder Plus (Figure 5) are solar-powered with large wings covered in solar cells. Because they carry solar cells and some batteries, they have tremendous altitude capability but relatively short range (daytime flights only) and small payloads. The Aerosonde falls into the “small” UAV class. It has a wingspan of only 2.9 meters and is powered by at 24cc engine that runs on unleaded gasoline fuel. The Aerosonde is launched from the rooftop of a speeding car, while the other UAVs require conventional runways.

History and UAV Limits The characteristics of UAVs are engineered to suit specific roles. UAVs originally evolved from target drones. The U.S. military extensively used the first UAVs during the 1960s in Vietnam [Wilson, 2003]. The Firebee and Compass Arrow programs developed these UAVs for reconnaissance, but the programs largely ended after the war. During the 1991 Gulf War, UAVs began to show their potential for flying in hostile conditions. UAVs are now extensively used by militaries around the world, and are being further developed by many countries. As with manned aircraft, there is a tradeoff between maximum altitude, payload, range, and airspeed. The lift provided by a wing is proportional to the area of the wing, the air density, and the square of the aircraft’s airspeed. Typically we measure the lift capability of an aircraft by the “wing loading.” Wing loading is the ratio of the weight of an aircraft to the area of the wings. A UAV with a low wing loading can fly at higher altitudes (lower density) with airspeed less than the speed of sound. A high altitude aircraft such as the Helios had a large wing and a low weight, giving it very small wing loading of 0.81 lbs./ft2. This low wing TABLE 1. Selective UAVs That Have Been Used or Proposed for Science Missions. UAV (manufacturer) Global Hawk (Northrop Grumman) Helios (Aeroenvironment) Pathfinder (Aeroenvironment) * Proteus (Scaled Composites) Altair (General Atomics) Altus-II (General Atomics) Aerosonde (Aerosonde)

Range (km) 25,000 200 200 5000 4200 5600 3000

Endurance (hrs) 36 15 14 14 32 24 30

Ceiling (km) 19.8 30.0 24.5 16.7 15.2 13.7 4.0

Payload (kg) 910 16 11 1000 300 150 5

Numbers are taken or calculated from manufacturers specifications. Actual performance will differ from these estimates. *Proteus is a manned aircraft that has only been used in a UAV demonstration.

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loading allowed the Helios to achieve altitudes near 100,000 feet. As payload weight is added, the wing loading increases and the aircraft can’t fly as high. In contrast to the Helios, the Global Hawk has a gross weight of 25,600 lbs and a wing area of 540 ft2, giving a wing loading of 47 lbs./ft2. The Global Hawk flies much faster than the Helios to gain the lift necessary to fly at 60,000 feet. The Global Hawk has an airspeed of about 343 kts at 20 km, while the Helios has an airspeed of about 150 kts at 30 km. Adding payload reduces the maximum altitude attainable by these aircraft. The drawback of a low wing loading comes from the necessity for a large wing. A UAV with large wing and a low weight is very susceptible to ground weather and cross winds. A takeoff becomes extremely difficult, since the aircraft wing acts as a large sail that can easily overturn the very light UAV. This creates the necessity for a very calm wind for both takeoff and landing. Airfields with extremely calm winds over significant periods are difficult to find.

Operations UAV operations are similar to operations for manned missions. However, UAVs require different command and control. The General Atomics MALE Altus aircraft (Figure 4) provides a useful example of operational constraints for UAVs. The Altus is a modestly size UAV with a length of about 7 m and a wingspan of 17 m. The Altus is powered by a gasoline engine, and can carry a payload of 150 kg to an altitude of about 14 km and has a range of about 5600 km. It has a C-band line-of-sight data link with a GPS/INS system. The Altus is easily disassembled into shipping cases, so it is easily transported and reassembled. The Altus is an excellent example of a versatile aircraft for atmospheric sampling. The Altus can be remotely piloted, or it can be run just like a satellite without human intervention. The waypoints are input to the program, and the aircraft executes those commands. The UAV can be communicated with using Ku-band

FIGURE 4. General Atomics Altus-I in flight while retracting landing gear in August 1997 (NASA DFRC, photographer Carla Thomas).

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(satellite) and C-band (microwave) communications to both direct and control the flight. While the Predator has a Ku-band link provides over the horizon communications, the Altus is not currently configured for over-the-horizon communications. The Altus has a forward-looking camera, such that flying the UAV is like sitting in an aircraft cockpit, but looking forward with narrow peripheral vision. Trailers containing the ground communications system with seating for scientists and the ground pilot are set up at the UAVs location, although the UAV can also be controlled from far distant locations because of the satellite link.

UAVs as Science Platforms UAVs in Field Experiments A few field campaigns have now been conducted with UAVs. In particular, the use of UAVs for Earth sciences and applications has been led by three programs: The Department of Energy’s Atmospheric Radiation Measurement (DOE/ARM) Program, the NASA Environmental Research Aircraft and Sensor Technology (ERAST) program, and the NASA UAV Science Demonstration Program (UAVSDP), and NASA’s Earth Sciences directorate. The DOE/ARM program has flown a number of UAV missions [Stephens et al., 2000]. Since 1993, DOE/ARM has sponsored 4 flight campaigns using 3 General Atomics UAVs: the Gnat-750, Altus-I, and Altus-II. The 1st science flight of a UAV was accomplished by the Gnat-750 in April 1994 when it made radiative flux measurements to an altitude of 7 km. They further flew a 26-hour duration flight over the CART site (5 Oct. 1996). In addition to these flights, the DOE/ARM program has developed several compact instruments suitable for UAV applications. The NASA ERAST program was started in 1994 to develop UAV technology and instrumentation. This program led to the development of a number of small instruments, and participated in the development of several high altitude aircraft. ERAST facilitated the development of (1) the Aurora Flight Sciences’ Perseus aircraft, which achieved a flight altitude of 60,260 feet in June 1998, (2) flights of the General Atomics Aeronautical Systems’ Altus II to 55,000 ft., with joint DOE/ ARM radiation experiments at Kauai, Hawaii, and wildfire imaging demonstration in September, 2001, and (3) flights of the solar-powered AeroVironment’s Pathfinder-Plus (Figure 5) and HELIOS, with a flight of the HELIOS to 96,863 feet from Kauai in August 2001. The NASA UAVSDP funded two science missions in 2001: the Altus Cumulus Electrification Study (ACES), and the UAV Coffee Project. The ACES objectives were to (1) validate a satellite instrument (the Lightning Imaging Sensor), (2) investigate lightning type, cloud-top optical energy, and power statistics, (3) lightning-storm relationships, and (4) storm electric budgets. The mission was flown in August 2002 from Key West, Florida. The Altus-II completed 13 flights and acquired 30 hours of data (http://aces.msfc.nasa.gov/index.html). The UAV

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FIGURE 5. AeroVironment’s Pathfinder solar-powered plane in flight on September 11, 1995 (NASA DFRC, photographer: Tony Landis).

coffee project was flown in September 2002 from Kauai using AeroVironment’s Pathfinder-Plus [Herwitz et al., 2004]. The project’s objective was to image the Kauai Coffee Company’s coffee farms for the detection of differences in coffee field ripening. The Pathfinder conducted a single 12-hour flight on September 30, 2002. The Pathfinder payload was an RGB digital camera and a narrow band digital multispectral camera. The Aerosonde has been used in a number of experiments (http://www.aerosonde. com/operations/). Two examples are the Fourth Convection And Moisture EXperiment (CAMEX-4), which was flown in the fall of 2001 and was sponsored by the NASA Earth Science Enterprise.. CAMEX-4 included 8 flights of the Aerosonde for studying the boundary at very low altitudes. The Aerosonde has also been flown from Barrow, Alaska, to study sea surface temperatures in the Arctic (Inoue and Curry, 2004). Use of UAVs for science in Europe has also been demonstrated with some flights of the EADS/IAI Eagle UAV from Kiruna, Sweden. The Eagle flights objectives were (1) give the Swedish military operational experience, (2) allow the Swedish Space Corporation to study UAVs as scientific platforms, (3) provide the Meteorological Institute (Univ. of Stockholm) UAVs for atmospheric sampling experience (Condensation Particle Counter). The flights of the Eagle were performed in June 2002 with 6 flights totaling 25 hours, that included one science flight of 5 hours (Hedlin and Abrahamsson, 2003).

UAV Problems and Current Uses UAVs present rather unique problems. First, at present they are unreliable. The General Atomics’ Predator RQ-1A had a mishap rate of 43 per 100,000 hours, and

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the RQ-1B had 31 [Peck, 2003]. This suggests that an Altus would be lost approximately once in every 3,000 flight hours. Such loss rates are 10 times larger than loss rates for current high-performance manned fighter aircraft. The NASA ERAST program was evaluating the Aurora Perseus B aircraft when it crashed on Oct. 1, 1999 (http://www.space.com/news/perseus_crash991004.html). The oneof-a-kind AeroVironment HELIOS crashed on June 26, 2003, and is not currently scheduled to be replaced. Such mishaps demonstrate that UAVs are still under development. While some UAVs are expendable, they are not randomly expendable. While inexpensive instruments on hazardous flights might be risked, test flights of one-of-a-kind multi-million dollar instruments cannot be risked. Second, aircraft must be flown frequently. The typical ER-2 or WB-57F can generally fly every day, but more typically fly every other day. For a typical 14day deployment, 5–12 flights totaling 40–100 hours can be flown. UAVs require more maintenance and can be more susceptible to weather conditions. During ACEs, the Altus-II was operational from August 2–30 (29 days), but only achieved a total of 30 hours of data. A substantial amount of these days were lost because of a problem with the specialized engines used by the Altus-II. The solar planes have extremely large wings, making them very susceptible to crosswind conditions on takeoff. During the UAV coffee project, AeroVironment’s Pathfinder-Plus had to scrub numerous times because of weather (http://geo.arc.nasa.gov/uav-nra/ index.html). Third, UAVs are not low cost. A UAV eliminates the necessity for a pilot and all of his life support, creating a weight dividend that can be used for fuel or payload. However, a UAV still requires complex communication with a ground system. This means that there is no “personnel” dividend. A UAV still requires ground support with a ground-pilot, mechanics, electricians, and computer specialists. Because UAVs are development aircraft, they typically have redundant command-and-control systems. This reduces the weight dividend that was gained by the absence of the pilot. Further, because UAVs are development planes, they have higher staffing than manned aircraft with established heritages. The cost of flying UAVs is higher than a manned aircraft because of (1) more personnel, (2) more equipment to operate the UAV, (3) UAVs are still under development, and (4) high insurance rates. Fourth, UAVs currently do not carry sufficient payload in comparison to manned aircraft. Much of the benefit of a manned aircraft is its ability to carry a number of heavy instruments. During the SOLVE-I campaign, the ER-2 carried 21 instruments, making it a functional chemical kinetics laboratory. The NASA WB-57F can carry 2300 kg to 19 km, and the NASA DC-8 can carry 13,600 kg to 12.5 km. As is shown in Table 1, the Global Hawk can carry a 910 kg payload to 19.8 km. The Global Hawk has the largest payload capacity of UAVs, but is considerably short of manned aircraft capacity. Finally, antennas and other equipment for redundant and robust aircraft control (e.g., over-the-horizon capabilities) consume valuable payload weight and volume. Fifth, UAVs are both very lightweight and typically built of composite materials. This creates problems for flying in severe weather conditions. For example, the crash of the Helios was attributed to turbulence in the lee of the island of

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Kauai. The turbulence caused the aircraft’s wing to abnormally bend beyond their breaking point. UAVs also cannot be flown where lightning strikes are probable, since lightning will melt or burn the composite material. Finally, UAVs have great difficulty gaining access to airspace. This is a severe restriction for Earth science. UAV flight request requires an extended period before the FAA approval. “Certified UAVs,” such as the Global Hawk, require only a few days, but are limited to take off and landing in restricted airspace. Nevertheless, atmospheric phenomena can move thousands of kilometers over the few days between the flight request and its approval. Good forecasting improves this, but errors of hundreds of kilometers would be expected. Many of these UAV problems will eventually be solved. However, it is clear that UAVs cannot currently replace manned aircraft. Developmental instruments that require a “hands-on” human operator are flown on manned aircraft such as the DC-8. This will probably never be done using a UAV. The UAVs currently being flown do not have the lift capability of aircraft such as the DC-8, WB-57F, or ER-2. There are no heavy lift UAVs in development. In spite of problems, NASA and other agencies continue to promote UAVs for Earth sciences. NASA and NOAA flew the General Atomics MALE Altair in the spring of 2005 for Earth science research probably out over the Pacific Ocean. NASA’s Research, Education and Applications Solutions Network (REASoN) has jointly developed a program with the U. S. Department of Agriculture to also use the Altair for monitoring forest fires in the Western United States. The spring 2005 NOAA and NASA mission flew the Altair from NASA Dryden Flight Research Center. The purpose of this mission was to test instruments and NOAA operational requirements. The Altair payload included remote and in-situ instruments for measurements of ocean color, and atmospheric composition and temperature; and a surface imaging system. In-situ composition measurements included ozone and long-lived gases such as halocarbons and nitrous oxide. The vertical distribution of water vapor has been remotely measured with passive microwave sensors. Six flights have been performed for a total of 53 hours of flight time. Flights reached altitudes up to 45,000 ft and have durations up to 20 hours. Flight objectives included sampling low-level jets in the eastern Pacific Ocean that bring moisture to the continental US; sampling regions of high potential vorticity at mid-latitudes that result from transport of polar air; and imaging of the Channel Islands National Marine Sanctuary (CINMS) to examine shorelines and evaluate the potential for marine enforcement surveillance. This NOAA/NASA experiment was a prelude to future routine operations that are targeted on improving both Earth science research, operational forecasting of both weather and pollution, and applications such as disaster assessment and mitigation.

Future Directions The future of UAVs in Earth science is bright, but the use of UAVs for Earth science is only in its infancy. The long-range of UAVs such as the Global Hawk and Altair bring phenomena within our reach that were unachievable in past

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decades. It is becoming possible to monitor the vast reaches of the Central Pacific or view Arctic Sea ice changes with a single aircraft launched from Hawaii or California. Slowly moving hurricanes, typhoons, and cyclones could be monitored and fully surveyed for tens of hours (surveillance), rather than with limited over flights by manned aircraft (reconnaissance). The extended duration of UAVs provides a new surveillance capability for Earth scientists. Most shortcomings of UAVs are being solved. Safety procedures, improved components and software, and continued military developments will increase the reliability of UAVs. New initiatives such as “Access 5” will address FAA flight restrictions and provide access to U.S. national airspace for UAVs. Routine access by Earth scientists to UAVs and access by UAVs to airspace will become common in the next few decades. UAVs bridge the temporal gap between the twice per day LEO satellite observations. UAVs also can provide the high vertical resolution and polar coverage that can’t be obtained by GEO satellite observations. The shorter endurance of manned aircraft makes them capable of performing reconnaissance (i.e., highresolution snapshots of phenomena) while the long endurance of UAVs makes them capable of remaining with phenomena and providing surveillance (i.e., highresolution continuous observations of a phenomena). In addition to their capability to fill a gap in the satellite coverage, UAVs provide the capability to enhance satellite remote sensing with high-resolution in-situ observations. These in-situ observations allow the satellite “radiance” to be built upward from forward model of radiative transfer. UAVs will evolve into “sub-orbital” satellite systems. In summary, manned aircraft and UAVs provide high-resolution observations that are (1) synergistic with satellite observations, and (2) fill observation gaps in satellite coverage. UAVs show the promise for major improvements in range and duration for Earth science. This improved capability moves UAVs from a reconnaissance role to a surveillance role. The improved capability also provides for a greatly enhanced capability of targeting phenomena. Scientists are using UAVs, and field missions have demonstrated the utility of UAVs. In the future, UAV missions will extend the scope of traditional manned-aircraft missions to those that are ultra-long, in dirty environments, and dangerous environments. In particular, these long-range, extended-duration UAVs will provide a logical extension of satellite missions to autonomous environmental monitoring as a part of sensor webs.

References Stephens, G.L., et al., The Department of Energy’s Atmospheric Radiation Measurement (ARM) Unmanned Aerospace Vehicle (UAV) program, Bull. Amer. Met. Soc., 81, 2915–2937, 2000. Hedlin, S., M. Abrahamsson, Demonstration of Eagle MALE UAV for scientific research, Proc. 18th Bristol International Conference on Unmanned Air Vehicle Systems, Dept. Aerospace Eng., University of Bristol, Bristol, United Kingdom, 31 March–2 April 2003. Herwitz, S.R., et al., Imaging from an unmanned aerial vehicle: Agricultural surveillance and decision support, Comp. and Electr in Agri., 44, 49–61, 2004.

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Inoue, J., A. Curry, Application of Aerosondes to high-resolution observations of sea surface temperature over Barrow Canyon, Geophys. Res. Lett., 31, DOI: 10.1029/2004GL020336, 2004. Peck, M., Pentagon Unhappy About Drone Aircraft Reliability, National Defense, 2003. Wilson, J.R., UAVS: A WORLDWIDE ROUNDUP, Aerospace America, 2003.

Chaper 9 U.S. Ground-based Campaign— PM Supersite Program KENNETH L. DEMERJIAN

Introduction The periodic assessment of new scientific information on the health and welfare effects of criteria air pollutants is required under the Clean Air Act and its amendments. As a result of new scientific information identified in its most recent particulate matter (PM) Criteria Assessment Document, EPA revised the PM National Ambient Air Quality Standard (NAAQS) in 1997 (U.S. EPA, 1997) Among the findings presented in this critical assessment, was the following. Overall, there is strong epidemiological evidence linking (a) short-term (hours, days) exposures to PM2.5 with cardiovascular and respiratory mortality and morbidity, and (b) long-term (years, decades) PM2.5 exposure with cardiovascular and lung cancer mortality and respiratory morbidity. The associations between PM2.5 and these various health endpoints are positive and often statistically significant. There is also extensive and convincing evidence for associations between short-term exposures to PM10 and both mortality and morbidity.

As a result of the PM2.5 (NAAQS) 1997 revision the U.S. Congress in response to public health concerns and in recognition of the uncertainty associated with some key aspects of the science of PM, decided in 1998 to provide funds for major research initiatives to address several policy-relevant scientific questions associated with PM2.5 regulation. Congress recognized that unlike gaseous pollutants, is not a single compound, but rather a complex mixture whose composition and morphology can vary in time and space and that these airborne particles have many sources and contain hundreds of inorganic and organic compounds. It also recognized that the new PM2.5 standard would placemany metropolitan areas in the U.S. into nonattainment status of the NAAQS. Figure 1 shows annual PM2.5 mass concentrations from 1007 locations in the U.S. for the period 1998–2000 [NARSTO, 2003]. The reds circles indicate locations exceeding the PM2.5 NAAQS. The yellow circles are locations likely to be violation in the near term, considering typical growth patterns. The significance of the magnitude of violations was another reason for Congressional interest.

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Annual PM2.5 Mass Concentration

Stations: AIRS FRM 1007, 1998-00 Canada TEOM 43, 1995-00 Canada Dichot 24, 1995-00

Annual PM2.5, µg/m3 > 15 10 - 15 < 10

FIGURE 1. Average PM2.5 concentrations in U.S and Canada (source Figure 6.7, NARTSO, 2003).

With these matters in mind, the Congress wanted the community to address two major PM science policy-relevant questions. Which component(s) of the PM mix is/are responsible for the statistical association with adverse health impacts and what are the mechanisms of action? What is the relationship between aerosol and aerosol precursor emissions and exposure of sensitive populations to the components identified above? As part of this research initiative a major workshop was convened involving the scientific and stakeholder communities. Its goal was to identify the key components and design parameters for a comprehensive measurement program to characterize ambient PM and important co-pollutants, in a way that optimizes information for multiple disciplines. A strategic ambient monitoring research program designed to develop, deploy, and evaluate measurement technologies for the monitoring the physical and chemical characteristics of particulate matter (PM) and its relationship to PM mass as measured by the Federal Reference Method (FRM). A competitive request for proposals resulted in the funding of series of programs. These measurement/research programs consisted of two Phase I and seven Phase II so-called PM Supersites, distributed geographically across the country (http://www.epa.gov/ttn/amtic/supersites.html) PM Supersites. Although quite diverse in their approaches, the “PM Supersite” programs shared several common themes across the country. These common themes were (1) to obtain atmospheric measurements to characterize PM, its constituents, precursors, co-pollutants, atmospheric transport, and source categories that effect the PM in the region, (2) obtain atmospheric measurements to support health effects and exposure research and to address the research questions and scientific uncertainties about PM source-receptor-exposure effects relationships, and (3) to

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conduct methods testing to compare and evaluate new PM measurement technologies for characterizing PM (physically and chemically). Each Supersite had specialized themes as well, including such activities as (1) characterizing PM urban/rural and summer/winter contrasts, (2) characterization PM diesel emission control technologies; PM organics speciation and transformations, (3) sources of new particle formation, (4) development and evaluation of deterministic and source apportionment modeling systems, (5) integrating and linking advanced PM size and chemical composition measurements with on going health studies, and (6) spatial characterization of ultrafine particles in the vicinity of major highways. The PM2.5 Supersite Phase I programs started in 1999, Phase II programs in 2000. The Phase I programs were typically 1–2 years in duration, while the Phase II programs operated over a 3–5 year period. Significant results have already been reported in the peer-reviewed literature (~147 papers) with many more in review and in preparation. Some specific results from the PM2.5 Technology Assessment and Characterization Study in New York (PMTACSNY), one of the seven Phase II PM Supersite programs are briefly discussed as a representative sample of the measurements/research performed under the program. The PMTACS-NY program to date has provided significant results addressing its three major objectives: ●

●

●

To measure the temporal and spatial distribution of the PM2.5/co-pollutant complex including photochemical oxidants and precursor species, aerosol size distribution and composition in urban and regional environments. To monitor the effectiveness of new heavy-duty vehicle emission control technologies [i.e., Compressed Natural Gas (CNG) and retrofit Diesel Filter— Continuously Regenerating Technology (DF-CRT)] introduced by the NYC Transit Authority for its MTA bus fleet and its potential impact on ambient air quality. To test and evaluate new measurement technologies and provide tech-transfer of demonstrated operationally robust technologies for routine network operation.

In the area of instrument evaluation, laboratory and field-based performance evaluations of water management system on continuous PM2.5 mass (R&P SESTEOM) monitoring systems [Schwab et al., 2003, 2004b, 2004c] have been performed. These studies have demonstrated that marked improvements have been made in reducing volatile losses of PM2.5 as compared to the standard 50 ° TEOM mass monitor. The studies also showed that mass differences are seasonal and likely associated with the content of NH4NO3 and semivolatile organic compounds in the PM. Based on these results recommendations regarding operational deploy of SES-TEOM systems and new generation Filter Dynamic Measurement System (FDMS TEOM) technology in routine PM2.5 monitoring networks have been made. In addition to the continuous mass monitors, semicontinuous PM2.5 sulfate and nitrate species measurement systems [Drewnick et al., 2003, Hogrefe et al., 2004, Hering et al., 2004] have been evaluated. These included advanced research measurement systems [Aerosol Mass Spectrometer (AMS), Particle-In-Liquid Sampler (PILS-IC)] as well as commercially viable

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systems [R&P 8400N/S and Continuous Ambient Sulfate Monitor (CASM)]. The CASM system is currently under development at Thermo, Inc. (under the designation 5020 Sulfate Particulate Analyzer). Over the course of these studies a variety of operational and maintenance issues have been identified and fed back to the instrument manufacturers, resulting in improvements in both the hardware and software of the instrumental systems. Overall the performance evaluations have been quite successful, demonstrating that the routine monitoring systems (R&P 8400S and 8400N; CASM) although having some systematic biases as compared to filter-based measurements can provide good-quality semicontinuous PM2.5 sulfate and nitrate concentration measurements. During the 2001 summer intensive field study chase studies on bus and truck emissions powered by Compressed Natural Gas (CNG), Diesel Filter Continuous Regenerating Technology (DF-CRT), and standard Heavy Duty diesel have been performed to characterize in-use emissions and the effectiveness of control technologies (CNG, DFCRT). The studies were performed in collaboration with NYC Transit Authority and NYS DEC and designed to specifically characterize the MTA bus fleet, which had introduced a fleet of CNG powered and DF-CRT retrofitted diesel buses. The study was to identify both the positive and potential negative impacts associated with the technologies. Results from these studies [Canagaratna, et al., 2004] showed significant reductions in in-use PM emissions from retrofitted DF-CRT buses, values quite consistent with those reported from chassis dynamometer studies. The studies also identified an NO2 slip problem, where NO2 used in the control process to react with carbon trapped on the oxidative catalytic surface is in excess under some load conditions. Although the net NOx emission remains unchanged, the fraction NO2/NOx can be as high as 50% under some conditions. In addition, significant benefits in the use of low-sulfur fuel (required by the DF-CRT technology and used for the entire MTA fleet) have been observed as well as major PM emission reduction in CNG powered buses [Shorter et al, 2005; Herndon et al., 2005]. A major negative impact observed from monitoring in-use CNG powered buses was the emission of formaldehyde. The emissions of SO2, H2CO, and CH4, measured using tunable infrared laser differential absorption spectroscopy, observed sulfur dioxide emissions from buses known to be burning ultra-low sulfur diesel (< 30 ppm(m) sulfur) were 16 times lower than buses burning normal commercial diesel fuel, nominally less than 300 ppm(m) sulfur. Formaldehyde levels observed in sampled in-use CNG bus exhaust plumes were in the 100s of ppb range, levels typically a factor ~15 times greater than their diesel-powered counterparts. Such ambient hot spot concentrations are unacceptable and the recommendation has been made to introduce an oxidative catalysis on these exhaust systems, to reduce formaldehyde emissions from CNG powered vehicles. Finally, substantial work has been performed characterizing the chemical composition particulate matter at urban and regional sites in New York State [Drewnick et al., 2004a, b; Schwab et al., 2004a]. A combination of short term field intensive studies and routine measurements have provided detailed insights as to the chemical composition of PM as a function of size and the seasonal variation of PM2.5

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composition across New York State. Figure 2 shows comparisons of mean PM composition from Aerosol Mass Spectrometer (AMS) measurements performed during PMTACS-NY 2001 Summer and 2004 Winter Intensive Field Studies. The results show somewhat higher PM sulfate and organic levels in the summer as compared to higher PM nitrate levels in the winter. Comparisons of summer versus winter AMS PM size distributions measurements (Figure 3) show a distinct difference in the mean mode aerosol diameter,

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FIGURE 3. Summer 2001 versus Winter 2004 comparison of PM size distribution in Queens, New York.

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with winter sizes smaller on average than summer sizes. We have evidence to suggest that this difference is likely due to secondary photochemical production processes involving gas to particle transformations of SO2 and VOCs which grow on existing aerosol surfaces. Also shown in Figure 3 is the disappearance of the bi-modal character of the PM organic in the wintertime data. The 70-nm mean mode diameter of PM organic in the summer-time data has been directly identified with primary diesel exhaust/lubricating oil emissions. A shifted in diesel exhaust emission size distribution with decreasing temperature has been observed by Abdul-Khalek et al., 1999, we believe that this and enhanced condensation on existing aerosol surfaces due to extreme cold weather conditions may explain these findings. We are currently studying this phenomenon in more detail as we reduce and analyze our aerosol sizing data. Finally, we have investigated the role of secondary PM production from local summertime photochemical transformations of SO2 to particulate sulfate and volatile organic compounds (VOCs) to secondary organic aerosol (SOA). We are particularly interested in an application of the AMS that looks extremely promising in providing estimates of SOA production and the role of VOC precursors in the attribution of PM organic mass in New York City. These results could have significant implications in the development of control strategies for both photochemical oxidants and PM. Using the monthly mean AMS data shown in Figure 4 we have estimate that 60% of the total summer organic (60% × 5.79 µg/m3) = 3.47 µg/m3 is attributable to “partially oxidized organics” (“POO”); while in the winter we estimate (35% × 3.85 µg/m3 = 1.35 µg/m3) is attributable to “POO” species. Organics based on AMS mass fragments [(m43& m57) and m44, for Prim and POO respectively]. We can demonstrate that photochemical production is negligible in the winter, thus the difference between summer and winter “POO” (3.47 − 1.35 = 2.12 µg/m3) provides an estimate of the summertime photochemical production of 2.12 µg/m3. Using measurements of OH concentration (Ren et al., 2003) and selected volatile organic compounds (NYS DEC PAMS) known to produce organic PM, we estimate the mean diurnal production over the course of the 2001 summer field intensive (see Figures 5a–c). The integration of the diurnal

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FIGURE 4. Summer 2001 versus Winter 2004 distribution of PM primary and partially oxidized.

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FIGURE 5A–C. Summer 2001 (a) mean diurnal OH measurements [based on Ren et al., 2003], (b) mean diurnal secondary PM organic production estimated from measurements of select VOCs and OH (c) mean diurnal integrated secondary PM organic production estimates in Queens, NY.

results gives a mean production of PM SOA for the campaign of 2.33µg/m3, a production rate very consistent with that estimated from the AMS organics mass fragmentation analysis. Results from the U.S. EPA “Supersite” Programs have provided compelling evidence for the key PM compositional components contributing to nonattainment of the PM2.5 National Ambient Air Quality Standard (NAAQS). These results indicate for example in New York City that carbon- and sulfur-based PM contributed more than 70% of the total PM mass throughout the year, with PM nitrate of some significance only during the cold season. In preparation of state implementation plans for PM, most cities will have to assess control strategies that deal with precursor emissions (SO2, VOC, and NOx) and primary emissions of OC/EC to achieve the PM2.4 NAAQS. Several pending national control programs slated for implementation in the next two years will likely be beneficial to the development of PM state implementation plans (SIPs). These included the new fuel sulfur rule and the 2007 heavy-duty diesel engine standards. But it is likely that additional controls will be needed. Preliminary results from Supersite field intensive studies in Queens, New York, suggest that an effective PM2.5 mitigation strategy would consider the implementation of a diesel filter trap catalytic converter retrofit control program for all heavy-duty diesel trucks operating in the major metropolitan areas have significant diesel truck traffic. Ultimately, whatever the strategy selected, it is critical to have measurements in place capable of tracking, through direct observation, changes in PM primary and precursor species concentrations in response to implemented control programs. This concept of Accountability in Air Quality Management Systems [Demerjian et al., 1995] is now accepted as a critical element in building credibility and public trust in our environmental regulations and their effectiveness in meeting air quality standards and anticipated improvements

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in environmental health. Continuation of “Supersite like” measurement programs is necessary to provide the essential monitoring data needed to track the progress of implemented air quality management strategies and addressing the reasons for their overall success or failure.

Reference Abdul-Khalek, I., D. Kittelson, F. Bear, The influence of dilution conditions on diesel exhaust particle size distribution measurements, Society of Automotive Engineers, Paper No.1999-01-1142. Albritton, D.L., D.S. Greenbaum, Atmospheric Observations: Helping Build the Scientific Basis for Decisions Related to Airborne Particulate Matter. Report of the PM Measurements Research Workshop, Chapel Hill, NC, 1998. Canagaratna, M.J., J.T. Jayne, D. Ghertner, S. Herndon, Q. Shi, J.L. Jimenez, P.J. Silva, P. Williams, T. Lanni, F. Drewnick, K. L. Demerjian, C. Kolb, D. Worsnop, Chase studies of particulate emissions from in-use New York City vehicles, Aerosol Science & Technology, 38, 555–573, 2004. Demerjian, K.L., A review of National Monitoring Networks in North America, Atmos. Environ, 34, 1861–1884, 2000. Demerjian, K.L., P.M. Roth, C. Blanchard. A new approach for demonstrating attainment of the ambient ozone standard, EPA/600/R-96/134 , U.S. EPA ORD Research Triangle Park, NC 27711, October 1995. Drewnick, F., J.J. Schwab, O. Hogrefe, S. Peters, L. Husain, D. Diamond, R. Weber, K.L. Demerjian, Intercomparison and evaluation of four semi-continuous PM2.5 sulfate instruments, Atmos Environ. 37, 3335–3350, 2003. Drewnick, F., J.J. Schwab, J.T. Jayne, M. Canagaratna, D.R. Worsnop, K.L. Demerjian, Measurement of ambient aerosol composition during the PMTACS-NY 2001 campaign using an aerosol mass spectrometer. Part I: Mass concentrations, Aerosol Science and Technology, 38(SI), 92–103, 2004a. Drewnick, F., J.T. Jayne, M. Canagaratna, D.R. Worsnop, K.L. Demerjian, Measurement of ambient aerosol composition during the PMTACS-NY 2001 campaign using an aerosol mass spectrometer. Part II: Chemically speciated mass distribution, Aerosol Science and Technology, 38(SI):104–117, 2004b. Dutkiewicz, V.A., S. Qureshi, A.R. Khan, V. Ferrara, J. Schwab, K. Demerjian, L. Husain, Sources of fine particulate sulfate in New York. Atmos. Environ., 38, 3179–3189, 2004. Hering, S., P.M. Fine, C. Sioutas, P.A. Jacques, J.L. Ambs, O. Hogrefe, K.L. Demerjian, Field assessment of the dynamics of the particulate nitrate vaporization using differential TEOM andautomated nitrate monitors, Atmos. Environ., 38, 5183–5192, 2004. Herndon, S.C., J.H. Shorter, M.S. Zahniser, J. Wormhoudt, D.D. Nelson, K.L. Demerjian, C.E. Kolb, Real-time measurements of SO2, H2CO and CH4 emissions from in-use curbside passenger buses in New York City using a chase vehicle” (in review ES&T). Hogrefe, O., F. Drewnick, G.G. Lala, J. J. Schwab, K.L. Demerjian, Development, operation and applications of an aerosol generation, calibration and research facility, new instruments and data inversion methods, Aerosol Science and Technology, 38(SI), 196–214, 2004. Hogrefe, O., J. Schwab, F. Drewnick, K. Rhoads, G.G. Lala, H.D. Felton, O.V. Rattigan, L. Husain, V.A. Dutkiewicz, S. Peters, K.L. Demerjian, Semi-continuous PM2.5 sulfate and nitrate measurements at an urban and a rural location in New York: PMTACS-NY Summer 2001 and 2002 campaigns., J. Air & Waste Manage. Assoc., 54, 1040–1060, 2004.

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Jayne, J.T. et al., Development of an Aerosol Mass Spectrometer for Size and Composition Analysis of Submicron Particles, Aerosol Sci. Technol., in press, 1999. Li, Z., P.K. Hopke, L. Husain, S. Qureshi, V.A. Dutkiewicz, J.J. Schwab, F. Drewnick, K.L. Demerjian, Sources of fine particle composition in New York City, Atmospheric Environment, 38, 6521–6529, 2004. Li, Y.Q., K.L. Demerjian, M.S. Zahniser, D.D. Nelson, J.B. McManus, S.C. Herndon, Measurements of formaldehyde, nitrogen dioxide, and sulfur dioxide at Whiteface Mountain using a dual tunable diode laser system, J. Geophys. Res. 109, D16S08, doi:10.1029/2003JD004091, 2004. NARSTO (2003). Particulate Matter Science for Policy Makers -A NARSTOAssessment, February 2003. http://www.cgenv.com/Narsto/. Ren, X., H. Harder, M. Martinez, R.L. Lesher, A. Oliger, T. Shirley, J. Adams, J. B. Simpas, W.H. Brune, J, HOx concentrations and OH reactivity observations in New York City during PMTACS-NY 2001, Atmospheric Environment, 37, 3627–3637, 2003. Ren, X., H. Harder, M. Martinez, R.L. Lesher, A. Oliger, J. B. Simpas, W.H. Brune, J. J. Schwab, K. L. Demerjian, Y. He, X. Zhou, H. Gao, OH and HO2 chemistry in the urban atmosphere of New York City, Atmospheric Environment, 37(26), 3627–3637, 2003. K.L. Demerjian, Y. He, X. Zhou, H. Gao, OH and HO2 chemistry in the urban atmosphere of New York City, Atmospheric Environment, 37, 3639–3651, 2003. Schwab, J.J., J. Spicer, H.D. Felton, J.A. Ambs, K.L. Demerjian, Long-term comparison of TEOM, SES TEOM and FRM measurements at rural and urban New York sites, in Symposium on Air Quality Measurements and Technology – 2002I, VIP-115-CD, ISBN 0-923204-50-4, Air and Waste Management Association, Pittsburgh, PA, 2003. Schwab, J.J., H.D. Felton, K.L. Demerjian, Aerosol chemical composition in New York state from integrated filter samples: Urban/rural and seasonal contrasts. J. Geophys. Res., 109, D16S05, doi:10.1029/2003JD004078, 2004a. Schwab, J.J., J. Spicer, K.L. Demerjian, J.L. Ambs, H.D. Felton, Long-term field characterization of TEOM and modified TEOM samplers in urban and rural New York State locations, J. Air & Waste Manage., 54, 1264–1280, 2004b. Schwab, J.J., O. Hogrefe, K.L. Demerjian, J.L. Ambs, Laboratory characterization of modified TEOM samplers, J. Air & Waste Manage., 54, 1254–1263, 2004c. Schwab, J.J., Y.-Q. Li, K.L. Demerjian, Semi-continuous formaldehyde measurements with a diffusion scrubber/liquid fluorescence analyzer. In Symposium on air quality measurement methods and technology—2004 [CD-ROM], Air and Waste Management Association, Pittsburgh, PA, 2004d. Shorter, J.H., S.C. Herndon, M. S. Zahniser, D. D. Nelson, J. Wormhoudt, K. L. Demerjian, C.E. Kolb, Real-time measurements of nitrogen oxide emissions from in-use New York City transit buses using a chase vehicle (in review ES&T). U.S. EPA 2004, The National Ambient Air Monitoring Strategy, Office of Air Quality Planning and Standards, Research Triangle Park, NC, April 2004. Yu, F., T. Lanni, B. Frank, Measurements of ion concentration in gasoline and diesel engine exhaust, Atmospheric Environment, 38, 10, 1417–1423, 2003. Zhou, X., H. Gao, Y. He, G. Hung, S.B. Bertman, K. Civerolo, J. Schwab, Nitric acid photolysis on surfaces in low-NOx environments: Significant atmospheric implications. Geophys. Res. Lett., 30, 2217, doi:10.1029/2003GL108620, 2003.

Part III Ground-Based Networks

Chapter 10 Ozone from Soundings: A Vital Element of Regional and Global Measurement Strategies ANNE M. THOMPSON

Abstract. Satellite instruments measuring ozone, predictive models of atmospheric trace gases, and models employing assimilation require well-resolved, accurate, and precise ozone profiles in the stratosphere and troposphere. These are provided by the ozonesonde instrument, typically an electrochemical concentration cell device flown on a balloon with a standard radiosonde. The design elements of a successful network of ozonesondes in the tropics (SHADOZ = Southern Hemisphere Additional Ozonesondes; Thompson et al., 2003a,b) are described. Sondes are a vital component of a sensor web as described by Schoeberl and Talabac [this volume] because profile variability can optimize retrieval accuracy as air mass types change. More generally, combined with assimilation, profiles from sondes and aircraft, as well as ground-based measurements are essential for accurate global measurement of tropospheric ozone.

Introduction Tropospheric ozone is of great interest locally as a pollutant. Ambient ozone concentrations greater than about 100 ppbv (parts-per-billion by volume) are deemed unhealthy and in many places emissions of the “precursor” gases (hydrocarbons, carbon monoxide, nitrogen oxides; for reactions, refer to Brune [this volume] leading to ozone formation are regulated. Ozone is measured continuously at the surface in hundreds of cities around the world. Globally, our concerns about ozone are its role in transboundary pollution, from city to city, region to region, and across continents and oceans. In addition, potential increases in the free tropospheric concentration of ozone are of interest because tropospheric ozone is a greenhouse gas. For global measurement, satellite instrumentation, aircraft, and sondes are employed. For example, GOME (acronyms given at end of text) and SCIAMACHY directly, and TOMS indirectly, are able to track ozone pollution with some resolution in the free troposphere. Aura, launched in July 2004, and presently in the testing phase, will expand tropospheric ozone measurements. Present satellite technology, however, does not resolve the vertical ozone distribution

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required for following trends [WMO, 1998] and for evaluating and constraining models. Thus, vertical profiling of ozone from aircraft [Cammas, this volume] and from balloon-borne ozonesondes, as presented in this chapter, are required. The roles of aircraft and sonde profiles in a sensor-web approach [Schoeberl and Talabac, this volume] to ozone measurement are complementary. Aircraft can operate regularly, taking frequent profiles with automated instrumentation over many locations, but within an altitude range limited to about 12 km [Marenco et al., 1998; Thouret et al., 1998]. Sondes are more labor-intensive, relatively expensive, and therefore are usually launched at weekly-monthly frequency [WMO, 1998] at a given site. However, in contrast to aircraft sampling, both tropospheric and stratospheric ozone (to 5–10 hPa) are sampled at 10–15 m resolution as the sonde ascends. Below, the rationale for ozone measurement by sondes and characteristics of ozone profiles are summarized. Then the design of a paradigmatic ozonesonde network for specific scientific objectives is described. Finally, the incorporation of ozonesondes in a sensor web that makes use of new satellite instruments and models is described.

Ozone Measurement Requirements and Profiles There are three major pollutants presently measured by satellite—ozone, carbon monoxide, and aerosols—all with lifetimes of days-months (Figure 2 of Brune, this volume, where PM 2.5 denotes aerosols). From aircraft, ground-based (e.g., lidar) and sounding observations, it is well-known that thin layers of the three pollutants can be very stable, making profiles highly variable. An example from an ozone sounding is shown in Figure 1 and the sonde instrument is shown in Figure 2. A set of tropical profiles from Ascension Island in the south Atlantic (8S, 14.4W), where ozone is characterized by low surface values (< 40 ppbv, purple trace at right) is illustrated. Aloft is pollution (layers with > 100 ppbv), and a distinct tropopause, signified by a rapid ozone increase and temperature inversion (red) at ~ 100 mb. Air parcel back-trajectories from pollution layers in the sounding show origins over Africa where biomass fires are active in September. See trajectories under “Ascension” and the corresponding year and date at: http://croc.gsfc.nasa.gov/shadoz The column-integrated total ozone amount for the Ascension sounding, with extrapolation based on the SBUV satellite profile climatology above 7 mb (59 DU, Dobson Units; one DU = 2.69 × 1016 molec/cm[2]), is 287 DU.

Design of a Tropical Ozonesonde Network and Key Findings The principles of an observing system incorporating ozonesondes are illustrated by a pair of networks we have designed. The discussion here is based on the SHADOZ (Southern Hemisphere Additional Ozonesondes) network we have organized in the southern tropics and subtropics. Since its inception in 1998, over

10. Ozone from Soundings Ascension Is, Lat. = −7.98 18 September, 2002 Long. = −14.42 19:34UT Temperature (C) −80 −60 −40 −40 0 20

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2300 ozone (and pressure-temperature-humidity, PTU) profiles have been archived at the SHADOZ Web site: http://croc.gsfc.nasa.gov/shadoz. The requirements and responsive design features of the network are summarized along with illustrations of scientific findings. In addition, we have found gaps in the present network and so we mention needs that new stations (and complementary approaches, e.g., aircraft profiling) can fulfill. The questions that SHADOZ addresses are as follows: ●

●

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Are tropospheric ozone satellite columns derived from satellite instruments (with TOMS, GOME, SCIAMACHY, and other instruments) accurate? Profile measurements are needed at sub-km resolution throughout troposphere and stratosphere. Only soundings meet this need with consistent sampling quality. What causes the zonal wave-one pattern observed in total equatorial ozone [Fishman and Larsen, 1987; Shiotani, 1992] and described further by Hudson and Thompson [1998] The geographic distribution of sounding sites must span the full tropical longitude range. What is the regional and temporal variability (over days-to-season scale) of tropical ozone? Are meteorological or photochemical factors primarily responsible

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FIGURE 2. Closeup of ozonesonde of electrochemical concentration cell (ECC) type.

for observed ozone variations? Weekly sampling is nominal; a few stations launch twice per month. Note that practical aspects of the ozonesonde measurement were factors determining SHADOZ site selection and sampling frequency. First, operational stations were chosen. Second, requirements of personnel and cost for ozonesonde operation (500–700 USD for instruments, balloon inflation gas and other supplies at each launch) and chemical preparation, launch and processing time (two days each) imposed practical limits. The SHADOZ network (as of 2004) is illustrated in Figure 3. Irene (late 1998), Malindi and Paramaribo (1999), Kuala Lumpur (data available in archive from 1998) joined after the 1998 initiation. Details of site location, operations and instrumentation appear in Thompson et al. [2003a; 2004]. Findings from SHADOZ, in response to the questions above, are as follows. ●

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Comparisons of total and tropospheric ozone from SHADOZ with TOMS satellite measurements appear in Thompson et al. [2003a,b] for 1998-2000. In general, TOMS total ozone is greater than the sonde value (example in Figure 4). In some cases, the sonde probably measures too low in the stratosphere [Smit and Sträter, 2004]. In other cases (as for Samoa, in Figure 4), the TOMS algorithm is known to overestimate tropospheric ozone (Figure 8 in Thompson et al., 2003b). A number of tropospheric ozone product comparisons have evolved for which SHADOZ comparisons are being made [Ziemke et al., 2002; Kim et al., 2004]. The wave-one feature that guided SHADOZ design, as it appears schematically in Figure 5, refers to a 15–20 DU higher total ozone column amount

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FIGURE 3. Map of SHADOZ stations current to 2004. See Thompson et al. [2003a, 2004] for latitude and longitude of each site. Operations at Tahiti ended after 1999. These stations began operating after the January 1998 start of SHADOZ: Irene (South Africa); Malindi (Kenya); Paramaribo (Surinam). SHADOZSite= American Samoa (145,171W) Integrated Sonde (@10mb), TOMS Overpass V8 Total Ozone Data

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FIGURE 4. Comparison of daily TOMS (version 8) ozone data from the satellite measuring as it passed over Samoa (14S, 171W), along with the Dobson total ozone measured from the surface at Samoa and total ozone integrated from the sondes launched in 1998–2001. Disagreement between the sonde and TOMS and the Dobson and TOMS (the satellite is typically higher in total ozone) is due to two factors: the satellite algorithm is artificially high in the troposphere; the particular sonde variety used in the measurement is apparently biased low in parts of the stratosphere [Thompson et al., 2003a; Smit and Sträter, 2004].

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−180

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FIGURE 5. Schematic representation of the wave-one feature in total ozone that is observed in the TOMS total ozone measurement and in total ozone integrated from SHADOZ sondes. Longitude is the horizontal axis. Exact location of maximum and minimum vary, as does the magnitude of highest ozone (5–20 Dobson Units).

over the Atlantic-Africa-western Indian Ocean region than over the midPacific. Analysis of the cross-sectional structure of ozone reveals that the integrated stratospheric column amount is invariant with longitude, to within the accuracy of the sonde technique. The tropospheric ozone column varies, however, with the same magnitude as found by the satellite measurement. Figure 6 (adapted from Thompson et al., 2003b, 2004), based on all the southern hemisphere SHADOZ data except for Irene, shows that more ozone tends

FIGURE 6. Cross-section of ozone mixing ratio based on 1998–2002 data averaged at 0.25 km intervals from the stations illustrated (except Irene and Paramaribo). More ozone is observed over the region spanned by eastern South America-Atlantic-Africa-western Indian Ocean than over the Pacific and eastern Indian Ocean. When integrated tropospheric ozone is compared between Atlantic and Pacific, the difference is equal to the typical “wave-one” magnitude. Thus, it is seen that the distribution of tropospheric ozone is responsible for the wave-one pattern. Stratospheric ozone, in contrast, is uniform over the same region (within experimental error; see Thompson et al., 2003a; Figures 12–14).

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to be concentrated in the free troposphere in the Atlantic-Africa-western Indian Ocean region. The causes of higher ozone in the region of the maximum include general circulation (subsidence indicated by downward arrows), a lower tropopause, and a greater concentration of ozone chemical precursors (from lightning and combustion) in the 40W to 60E region, compared to the Pacific [Oltmans et al., 2001]. Analysis of temporal variability at a single SHADOZ station (e.g., Figure 9 in Thompson et al., 2003b) shows signals with a 3–7-day period. For example, this type of periodicity has been deduced from Nairobi record (Figure 7). Unexpected regional variability emerged in several studies. For example, in early 1999, during the Aerosols99 trans-Atlantic ship crossing [Thompson et al., 2000] and the time of the INDOEX experiment [Chatfield et al., 2004], higher ozone in the troposphere appeared over the southern tropical Atlantic than at the same latitudes north of the ITCZ where biomass burning was active. This phenomenon is referred to as the “tropical Atlantic paradox” (depicted in ozone latitudinal mixing ratio cross-section in Figure 8). When ozone seasonality is compared at Paramaribo (Figure 8) to that for a southern hemisphere SHADOZ site (see Figure 8 in Thompson, 2004), there is a contrast, primarily in the January-to-March period. This is due to different “seasonality” at Surinam in terms of ITCZ position and precipitation and winds compared to the southern tropical stations [Peters et al., 2004]. Nairobi 1998-2001 Data - 0.25km Bins

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FIGURE 7. Ozone mixing ratios, 0.25-km means, from weekly soundings over Nairobi from 1998–2001 (adapted from Thompson et al., 2004). Time-series analysis shows cycles on an annual basis, semi-annual, 30–60 days (the frequency of the Madden–Julian Oscillation) and a signal less than 10 days. Advective and convective processes, introducing cleaner or more polluted air masses to the Nairobi area, are responsible for much of the short-term variability [Chatfield et al., 2004].

138

Anne M. Thompson Paramaribo, Surinam (6N, 55W) Ozone Mixing Ratio [ppbv]

Altitude [km]

15

10

5

0 Jan

Feb

Mar

Apr

May

Jun

Jul

Aug

Sep

Oct

Nov

Dec

Monthly X-sections including years 1998-2002 Aerosols99 Ozone Mixing Ratio (ppbv) Altitude (km) vs. Latitude (deg) Zones: 1

2

ITCZ

3

4

15

10

5

0 30

20

10

−10

0

−30

150

100

85

75

65

55

45

35

25

15

5

ozone (ppbv)

−20

FIGURE 8. (A) Latitudinal ozone cross-section below 17 km measured by soundings during the Aerosols99 trans-Atlantic cruise (figure adapted from [Thompson et al., 2000]). The higher ozone in the lower and middle troposphere over the north tropical Atlantic (designated “2”) in the figure is due to pollution from biomass fires over north Africa. However, in zone “3”, south of the equator where the wet season is prevalent, integrated tropospheric ozone is higher, due to the concentrated higher-ozone amounts in the middleupper troposphere. (B) Seasonal variation in tropospheric ozone over Paramaribo, based on soundings averaged over each month during 1999–2003.

10. Ozone from Soundings ●

139

Another intriguing variability in ozone regional distributions has been described by Sauvage et al. [2004]. The ozone distribution based on the MOZAIC [Marenco et al., 1998; Thouret et al., 1998] profiles from aircraft landings and takeoffs over Africa during December-January-February shows a latitudinal distribution without the “paradox” feature. As one would expect from proximity to biomass fires, there is higher ozone in the lower troposphere north of the equator (integrating to greater total tropospheric ozone) than there is in the southern hemisphere. Relatively enhanced upper tropospheric ozone (similar to that shown in Figure 8) is observed in the aircraft profiles over the African continent, similar to the Atlantic pattern, and consistent with the largescale feature of subsidence depicted in Figure 6.

These discoveries of contrast between northern and southern hemispheric ozone and the apparent lack of a “paradox” over the African continent are leading to a revised design for SHADOZ as we move to the Aura era [Bhartia, this volume]. Potential new regions for SHADOZ sites are in west Africa, southern Asia, the middle East, central America. In addition, interferences of aerosols with ozone retrieval cause us to consider more pollution as a criterion for new network locations.

The Ozonesonde as an Essential Element of Integrated Atmospheric Chemical Observations Figure 9 shows schematically that ozonesondes can integrate observations and assimilation model representations of ozone distribution in the atmosphere. Satellites are not yet able to resolve ozone concentrations in the vertical to the

Satellite

SONDE

Models, Assimilation

Ground-based, Aircraft

FIGURE 9. Schematic of sondes as an integrating element of information among models, aircraft profiles, satellite column measurements for the derivation of tropospheric ozone. Neither models nor satellite are able to consistently derive the location of the tropopause and aircraft profiling encompasses this boundary only when it is below about 200 hPa. Sondes provide locate the tropopause region with the resolution required for guiding satellite algorithms and corresponding assimilation.

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extent required for tracking pollution transport. Nor do they locate the tropopause reliably enough for separation of stratospheric and tropospheric ozone [Thompson et al., 2001]. Sondes can do this consistently but their geographical and temporal coverage is limited. Thus, multiple observing elements need to be present in order to obtain regional and global three-dimensional ozone distributions. There are several ways that sondes can be used to optimize ozone retrieval from satellite. One is that buv instruments (e.g., TOMS, OMI) start with first-guess profiles in which stratospheric ozone is varied according to total ozone amount and latitude. Ozonesonde climatologies like SHADOZ have shown that latitude alone (and season) are inadequate for specifying a first-guess profile (resulting in the disagreement shown in Figure 4). However, sonde climatologies can be prepared according to total ozone [Hudson et al., 2003] with statistically derived corresponding stratospheric and tropospheric ozone profiles. Models have evolved to give quite accurate tropopause forecasts. This tropopause information can be combined with total ozone and first-guess tropospheric and stratospheric profiles for a given region and tropopause location to derive improved ozone column and profile amounts. Another approach that would make use of sondes is to combine them with multiple satellite information and cloud and aerosol data. The latter can originate from aerosol and cloud sensors aboard Aqua and correlative ground-based instrumentation. Possibilities include MODIS, refined with observations from AERONET and MPL-Net, and aerosol models. Acknowledgements The author thanks J. C. Witte (SSAI at NASA Goddard Space Flight Center) for graphical analysis and J. B. Stone (Penn State University) for assistance with the manuscript. The author is grateful to W. H. Brune (Penn State University) and P. Di Carlo (University of L’Aquila) for the opportunity to participate in ISSAOS, September 2004.

Acronyms AERONET = Aerosol Robotic Network (http://aeronet.gsfc.nasa.gov) AIRS = Atmospheric Infrared Sounder (on Aqua, 2002-) GOME = Global Ozone Monitoring Experiment (on ERS-1, 1995-) INDOEX = Indian Ocean Experiment (1999) ISSAOS = International Summer School on Atmospheric and Oceanic Science MODIS = Moderate-resolution Imaging Spectrometer (on Terra, 1999-; on Aqua, 2002) MOZAIC = Measurements of Ozone from Airbus In-service Aircraft (1994-) MPL-Net = Micropulse Lidar Network (http://mplnet.gsfc.nasa.gov) OMI = Ozone Monitoring Instrument (on Aura, 2004-) SCIAMACHY = Scanning Imaging Absorption Spectro-Meter for Atmospheric Charto-graphY SHADOZ = Southern Hemisphere Additional Ozonesondes TES = Tropospheric Emission Spectrometer (on Aura, 2004-) TOMS = Total Ozone Mapping Spectrometer (on Earth-Probe, 1996-)

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References Brune, W.H., this volume, 2006. Cammas, J-P., this volume, 2006. Chatfield, R.B., H. Guan, A.M. Thompson, J.C. Witte, Convective lofting links Indian Ocean air pollution to paradoxical south Atlantic ozone maxima, Geophys. Res. Lett., 31, L06103, doi: 10.129/2003GL018866, 2004. Fishman, J., J.C. Larsen, Distribution of total ozone and stratospheric ozone in the tropics -Implications for the distribution of tropospheric ozone, J. Geophys. Res., 92, 6627–6634, 1987. Hudson, R.D., A.D. Frolov, M.F. Andrade, M.B. Follette, The total ozone field separated into meteorological regimes. Part I: Defining the regimes, J. Geophys. Res., 60(14), 1669–1677, 2003. Hudson, R.D., A.M. Thompson, Tropical Tropospheric Ozone (TTO) maps from TOMS by a modified residual method, J. Geophys. Res., 103, 22,129–22,145, 1998. Kim, J.H., S. Na, M. Newchurch, K.J. Ha, Comparison of scan-angle method and convective cloud differential method in retrieving tropospheric ozone from TOMS, Environmental Monitoring and Assessment, 92, 25–33, 2004. IGOS Integrated Global Observing Strategy and the IGOS partnership (report) (http://ioc.unesco.org/igospartners/). Marenco, A., V. Thouret, P. Nedelec, H. Smit, M. Helten, D. Kley, F. Karcher, P. Simon, K. Law, J. Pyle, G. Poschmann, R. Von Wrede, C. Hume, T. Cook, Measurements of ozone and water vapor by Airbus in-service aircraft: The MOZAIC airborne program, An overview, J. Geophys. Res, 103(D19), 25631–25642, 10.1029/98JD00977, 1998. Oltmans, S.J. et al., Ozone in the Pacific tropical troposphere from ozonesonde observations, J. Geophys. Res., 106, 32503–32526, 2001. Peters, W., M.C. Krol, J.P.F. Fortuin, H.M. Kelder, C.R. Becker, A.M. Thompson, J. Lelieveld, P. J. Crutzen, Tropospheric ozone over a tropical Atlantic station in the northern hemisphere: Paramaribo, Surinam (6N, 55W), Tellus B, 56, 21–34, 2004. Randriambelo, T., J-L. Baray, S. Baldy, A.M. Thompson, S.J. Oltmans, P. Keckhut, Investigation of the short-term variability of tropical tropospheric ozone, Annales Geophysiques, 21, 2095–2106, 2003. Sauvage, B., V. Thouret, J-P. Cammas, F. Gueusi, G. Athier and P. Nédélec, Tropospheric ozone over Equatorial Africa: Regional aspects from the MOZAIC data, Atmos. Chem. Phys. Disc., SRef-ID: 1680-7375/acpd/2004-4-3285, 2004. Schoeberl, M.R. and Talabec, this volume, 2006. Shiotani, M. Annual, quasi-biennial and El Nino-Southern Oscillation (ENSO) time-scale variations in Equatorial total ozone, J. Geophys. Res., 97, 7625–7634, 1992. Smit, H.G.J., W. Sträter, JOSIE-1998: Performance of ECC Ozone Sondes of SPC-6A and ENSCI-A Type, in WMO Global Atmospheric Watch Report Series, No. 157 (Technical Document No. 1218), World Meteorological Organization, Geneva, 2004a. Smit, H.G.J., W. Sträter, JOSIE-2000: The 2000 WMO international intercomparison of operating procedures for ECC-sondes at the environmental simulation facility at Jülich, in WMO Global Atmospheric Watch Report Series (Technical Document), WMO Global Atmosphere Watch report series, No. 157 (Technical Document No. 1218), World Meteorological Organization, Geneva, 2004b. Thompson, A.M., et al. A tropical Atlantic ozone paradox: Shipboard and satellite views of a tropospheric ozone maximum and wave-one in January-February 1999, Geophys. Res., Lett., 27, 3317–3320, 2000.

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Thompson, A.M., R.D. Hudson, A.D. Frolov, J.C. Witte, T.L. Kucsera, Tropospheric ozone from space: Tracking pollution with the TOMS (Total Ozone Mapping Spectrometer) instrument, IGARSS Proceedings, Sydney, Australia, Meeting, July 2001, IEEE Publ., Piscataway, NJ, 2001. Thompson, A.M., et al., Southern Hemisphere Additional Ozonesondes (SHADOZ) 1998–2000 tropical ozone climatology: 1. Comparison with TOMS and ground-based measurements, J. Geophys. Res., 108, 8238, doi: 10.129/2001JD000967, 2003a. Thompson, A.M., et al., Southern Hemisphere Additional Ozonesondes (SHADOZ) 19982000 tropical ozone climatology. 2. Tropospheric variability and the zonal wave-one, J. Geophys. Res., 108, 8241, doi: 10.129/2002JD002241, 2003b. Thompson, A.M., J.C. Witte, S.J. Oltmans, F.J. Schmidlin, SHADOZ (Southern Hemisphere Additional Ozonesondes): A tropical ozonesonde-radiosonde network for the atmospheric community, Bull. Am. Meteorol. Soc., 85, 1549–1564, 2004. Thouret, V., A. Marenco, J.A. Logan, P. Nedelec, C. Grouhel, Comparisons of ozone measurements from the MOZAIC airborne program and the ozone sounding network at eight locations, J Geophys. Res., 103(D19), 25685–25720, 10.1029/98JD02243, 1998. WMO (World Meteorological Organization), SPARC/IOC/GAW Assessment of Trends in the Vertical Distribution of Ozone, ed. By N. Harris, R. Hudson and C. Phillips, SPARC Report No. 1, WMO Global Ozone Research and Monitoring Project, Report No. 43., 1998. Ziemke, J.R., S. Chandra, P.K. Bhartia, “Cloud Slicing”: A new technique to derive upper tropospheric ozone from satellite measurements, J. Geophys. Res., 106 (D9), 9853–9868, 10.1029/2000JD900768, 2001.

Chapter 11 Lidar Networks VINCENZO RIZI

Introduction This lecture covers some aspects of the lidar networks as systems for the observation of the atmosphere. After a short discussion of the lidar technique and of the typical lidar products, a few words are devoted to a partial list of the recent lidar network experiments. Finally, a discussion of the scientific problems and studies that can be addressed within a coordinated lidar network is presented (the peculiar example is a Raman-based lidar network, EARLINET, supported in 2000-2003 period by European Community). The LIDAR (LIght Detection And Ranging) technique is quite diffuse among atmospheric studies. Figure 1 shows an elementary representation of this remote probing and sampling (with light as information carrier) of the atmosphere. The laser emitted photons, while traveling in the atmosphere, are scattered, by different processes, of different efficiencies, involving the laser light features and the different components of the atmosphere. The backscattered photons are collected by a telescope and detected by a configuration of photo-detectors as photomultipliers, etc.. Collecting the backscattered photons (not necessary at the same wavelength of the emitted ones) as a function of their flight time, it is possible to discriminate the photons returning from different ranges/altitudes. The number of collected photons depends of the atmospheric optical transmission (along the upward and downward travel), which origininates from the scattering and absorption processes, and of the local backscattering process. m m m m L (s) ? L o $ T up (s) $ 7 backscattering (s) A $ T down (s) d4X (1) r Ll (s) is the number of collected photons from the range s; Lol is the number of photons emitted by the laser; [backscattering(s)] includes all the effective scattering processes (i.e., Rayleigh–Mie scattering, Raman scattering, resonant scattering) that return photons into the receiver; Tuplo (s) and T ldown (s) are the atmospheric transmission function along the travel forth and back from s range, and dΩ/4p is 2 the solid angle subtended the receiver (? 1/s ). o

o

o

143

144

Vincenzo Rizi

LIDAR

backscattering

Laser emitted photons

Backscattered photons Rayleigh/Mie/Raman

Laser

Telescope

DAQ

FIGURE 1. Simple representation of lidar sampling of the atmosphere, break DAQ: data acquisition system.

The baseline design of a lidar system consists of a laser, which transmits short light pulses in the atmosphere, and a receiver, usually a telescope and a detector, which collects the backscattered returns as a function of the (light pulse) time of flight (i.e., range from lidar). Quite diffuse commercial interference filters and dichroic beam splitters allow us to design a spectrometer-wise receiver: it is possible to discriminate the lidar returns according to the wavelengths.

Standard Lidar Systems This paragraph reports a short review of the standard lidar applications devoted to detect the atmospheric composition and structure. The different atmospheric components can contribute to the backscattering and transmission terms into the lidar equation. Figure 2 illustrates various lidar applications.

Aerosol Raman Lidar A relatively new and powerful lidar technique is based on the Raman backscattering of some molecular components of the atmosphere. This paragraph is dedicated to a simple outline of a typical aerosol Raman lidar system

11. Lidar Networks

Aerosol lidar

i.e., stratospheric aerosols

Transmitter

145

Differential Absorption lidar Transmitter

i.e., stratospheric ozone

Receiver Return Pulse Power

Receiver Return Pulse Power

Signal Distance

backscattering increase Aerosol backscattering

relatively more attenuation

λ0

absorption cross section Reference

σ1 σ0

Signal λ2

λ1 λ0

λ

Distance

Ratio

Distance

Concentration

Distance Distance

Raman aerosol lidar

i.e., tropospheric aerosol

Raman water vappor lidar Transmitter

i.e., tropospheric water vapor

Transmitter Receiver Receiver Return Pulse Power

more backscattering more attenuation no backscattering

Rayleigh/Mie signal λ0 N2 Raman/anel astic signal λ0+∆λN2 Distance

aerosol backscattering aerosol transmission

Distance

N2 backscattering

Return Pulse Power

H2O backscattering H2O mixing ratio

N2 Raman/anel astic signal H2O Raman/anel astic signal λ0+∆λN2 λ0+∆λH2O Distance

Distance

Distance

FIGURE 2. Aerosol lidar (top-left panel): an increase in the aerosol concentration reflects directly into the backscattering term of lidar return. Differential absorption lidar (top-right panel): this system makes use of two source wavelengths, one is strongly absorbed by the molecular specie to be measured (i.e., O3), the other shows a corresponding negligible absorption; the two lidar signals differ in the transmission part, from this difference the concentration of the molecular specie can be retrieved. Raman aerosol lidar (bottom-left panel): it is one source system, but the detection of lidar returns is performed also at a different wavelength according to the Raman inelastic scattering (i.e., from N2 or O2 Raman scattering, more details are discussed in the next paragraphs); the Raman lidar signal is affected by the aerosol presence only in the transmission part. Raman water vapor lidar (bottom-right panel): the lidar collects signals at specific Raman shifted wavelengths (i.e., N2 and H2O vapor), in this case the transmission part of the lidar return have the same behavior, as a consequence, a direct comparison of the Raman backscattered signals and a calibration procedure allow us to estimate the relative fraction of H2O.

The molecular scattering in the atmosphere consists of Rayleigh (elastic) scattering and rotovibrational Raman (anelastic) scattering. What is called Rayleigh scattering consists of rotational Raman lines and the central Cabannes line. In Raman scattering the cross section is about 3 orders of magnitude smaller than the corresponding Rayleigh cross section, and the scattered signal consists of radiation that has suffered a frequency shift that is characteristic of the stationary energy states of the irradiated molecule.

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The Raman spectroscopy enables a trace constituent to be both identified and quantified relative to the major component of a mixture. As an example: the nitrogen and oxygen molecules (whose atmospheric mixing ratios are constant and known) show Raman shifts (vibrational-rotational transitions) of about 2327 cm−1 and 1556 cm−1, respectively. Focusing on the lidar sounding of high-resolution profiles of atmospheric aerosols; the advantages offered by the combined Raman–Rayleigh lidar detection are quite evident: more information about the measured portion of the atmosphere is available. Some details are reported to make clear this point: with the same formalism used in the previous section, the elastic/Rayleigh lidar signal has the following form: mo

mo

mo

mo

L (s) = L o $ T mol (s) $ T aer (s) $ = v mmol (r) $ n mol (s) + o

3

# drrr

2

0

Q bck (r, m, m o ) n aer (s, r) G $

(2)

m m $ d4X r $ cx L $ T mol (s) $ T aer (s) o

o

where Lol

number of photons emitted by the laser;

smollo (p)

Rayleigh backscattering cross section:

o

P (r) m o v mol , P (r) is the backscattering phase function; 4r Qbck (r,m,lo) Mie backscattering efficiency; mo v mol (r) =

Tlmol (s) molecular (Rayleigh + vibrational Raman) scattering transmission at wavelength λ; Taerl (s) aerosol scattering transmission; dΩ solid angle subtended by receiver; tL laser pulse duration (c is the speed of light); and the atmospheric transmission functions have the following main dependence from the atmospheric components: T mol (s) = exp e- # v mmol n mol (s) ds o s

m

(3)

0

T aer (s) = exp f- # = # drrr 2 Q ext (r, m, m) n aer (s, r) G ds p s

m

0

3

(4)

0

nmol atmospheric molecular number density; naer aerosol size distribution: number density of aerosol with radius between r and r+dr; smoll Rayleigh total cross section; Qext (r,m,l) Mie extinction efficiency of an aerosol of radius r, and refractive index m.

signal * (range)2 (a.u.)

11. Lidar Networks 1.00

147

Rayleigh/elastic air and aerosols

0.10 cloud 0.01 0

1

2

3 4 range (km)

5

6

7

FIGURE 3. Example of elastic signal of a vertical lidar sounding the troposphere in presence of a typical low-level cloud. This profile has been collected on October 8, 2001, with the UV Raman/Rayleigh lidar at L’Aquila, Italy. The signal is the summation of about 54,000 laser shots (about 30 min); the vertical resolution is 30 m.

An example of Rayleigh lidar return is shown in Figure 3. These data have been taken in presence of a low-level cloud, within the cloud range the signal enhances due to the increasing of the effective aerosol backscattering. The anelastic/Raman lidar signal of the ith molecule is Lm = m i

o+

Dm i

m (s) = Lmo $ T mol (s) $ T aerm (s) $ 8 v mRaman (r) $ n i (s) B $ o

o

o

(5)

i

dX $ cx $ T m (s) $ T m (s) L mol aer 4r i

with ∆li

i

is the Raman shift of the ith molecule

R V -1 S W (Do N . 2327cm - 1 ; Do O . 1556cm - 1, Dm N = S c 1 - Do N m - m o W; S mo W T X R V -1 S W Dm O = S c 1 - Do O m - m o W S mo W T X 2

2

l

2

2

2

2

i s Raman (p) is the Raman backscattering cross section of the ith molecule: e.g., for N2: ≈3.5×10−30 cm2 sr−1, for O2: ≈ 4.6×10−30 cm2sr−1 at l = 337 nm. Figure 4 shows the N2 Raman lidar signal corresponding to the measurement of Figure 3, the cloud presence appears as an attenuation of the lidar return. Comparing Figures 3 and 4, it is evident that in the elastic channel the backscattering contribution of the cloud dominates, while in the Raman anelastic channel the main feature is the extinction in the cloud range. In general, the information retrieved by coupling Rayleigh and Raman signals are complementary and might allow a quicker estimation of the atmospheric transmissivity. It should be noted that in the elastic channel the aerosol scattering contribution is present in both the transmission and backscattering components, while in the

Vincenzo Rizi

signal * (range)2 (a.u.)

148

1.00 Raman/anelastic N2 cloud transmission 0.10 cloud

0

1

2

3 4 range (km)

5

6

7

FIGURE 4. As Figure 3, but for the anelastic/Raman signal of nitrogen. Note the attenuation of the Raman signal in the cloud range.

Raman channels it appears only in the transmission. Therefore, quantitative measurements of aerosol optical properties using a lidar system that measures only aerosol elastic backscatter requires accurate system calibration and assumptions regarding aerosol optical properties (i.e., relationship between the aerosol backscatter and transmission). A combined Raman–Rayleigh lidar allows us to greatly reduce the dependence of the reconstructed transmission function on calibration and assumptions, since both aerosol transmission and backscatter are independently and directly measured.

Aerosol Characterization by Lidar A short discussion of the performances of Raman lidar in retrieving the aerosol optical properties follows in the next sections. It is assumed that the atmospheric molecular number density is known, as well as the details of the Rayleigh scattering. To be more synthetic, we define the following quantities widely used within the lidar community: 3

b maer (s) = o

# drrr

2

Q bck (r, m, m o ) n aer (s, r)

(6)

2

Q ext (r, m, m o ) n aer (s, r)

(7)

0 3

a maer (s) = o

# drrr 0

l l baer (s) is the aerosol volume backscatter coefficient, and aaer (s) is the aerosol extinction coefficient. o

o

Rayleigh Lidar l If only the elastic signal is available, to estimate the aerosol transmission, Taer (s), one must assume the so-called lidar ratio, LR in Eq. (8), which is the relationship between the aerosol extinction and backscatter: a m (s) LR = aer (8) b maer (s) o

o

o

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149

This makes it possible to solve Eq. (2) for the aerosol extinction. The obtained expression is the standard Bernoulli equation that can be integrated (Klett method), knowing a calibration height, where the boundary value of aerosol extinction can be estimated. This inversion method suffers from the fact that two physical quantities, the aerosol backscatter and the aerosol extinction, must be determined from only one measured lidar signal. Depending on the aerosol type, the lidar ratio can vary in a wide range, and in practice, no information on the required height profile of the LR is usually available; on the other hand, if the best available constant lidar ratio is used, large uncertainties in the aerosol extinction profile remain due to the height dependence of LR. Anelastic/Raman Lidar (Additional Detection of N2 Raman Scattering) In this case, the anelastic (due to molecular nitrogen Raman scattering) lidar sigl nal is known [Eq. (5)]. To derive Taer (s), in addition to the calibration value, the wavelength dependence of the aerosol extinction should be known. This is done assuming the exponent, k: o

log f k=

log (T aerm (z)) p log (T aerm (z)) o

N2 i

(9)

log d m o n mN 2

Equation (5) can be solved for Taerlo (s): log ` T aerm (s)j = o

1 log 1 + (m o /m N ) k 2

R V S W Lm (s) m 1 + _ m /m i E ; S W [T mol (s)] (10) m m SS L a $ v Raman (r) $ n N (s) $ dX WW 4r T X For aerosol particles and water droplets with diameters comparable with the measurement wavelength k = 1 is appropriate. 1

N2

n

n

n

N2

N2

2

An Example with Real Lidar Signals This example refers to a real standard session of measurements (UV Rayleigh/Raman lidar located at L’Aquila, Italy, 42.35N, 13.22E, site elevation 683m a.s.l.) during a clear-sky night. Figure 5 shows the lidar signals. Coupling the analysis of the Rayleigh/elastic and Raman/anelastic signals, it is possible to retrieve the aerosol extinction and backscatter coefficients, and the lidar ratio: these measurements are shown in Figure 6. A sketch of the procedure is the following (the atmosphere density is known): ●

the extinction is derived from the anelastic signal assuming the scaling with wavelength (k = 1);